oversize

Guidebook to

HE GEOLOGY OF\ORTHEASTERN MAINE
AND NEIGHBORING NEW BRUNSWICK

' i

•

David C. Roy
Boston College

Edited by

and

Richard S. Naylor
Northeastern University

Contributors

* Robert A. Ayu$o
Harold W. Borhs
Thomas Brewer
D. W. Caldwell
William H. Forbes
Claude Gauthier
Andrew 1ST. Genes
Terence Hamilton-Smith
Rudolph Hon
Andrew E. Kasper, Jr.
J. Steven Kite N
Glen.G. Lutes

Richard S. f^aylorY
Robert B. Neumans
William A. Newman
Philip H Osberg
m. K, Pickerill j
Douglas W. Rank^i
N. Rast \F?

David C. Roy/f
C. St. Peter/
Jacque^fhibault
David R. AA/ones

The 72nd Annual lyieeting of the
New England lntercollegiaj£ Geological Conference

■

*

'

University of
new Hampshire

Library

Copies of this guidebook may be obtained for $10.00
each (U.S. funds) by sending requests, prepaid, to:

David C. Roy

Department of Geology and Geophysics

Boston College

Chestnut Hill, MA

02167

Printed by Boston College Press

1980

UNH LIBRARY

\S

MbOD 003D3 TMMO

DISCLAIMER

Before visiting any of the sites described in the New
England Intercollegiate Geological Conference
guidebooks, you must obtain permission from the
current landowners.

Landowners only granted permission to visit these sites
to the organizers of the original trips for the designated
dates of the conference. It is your responsibility to obtain
permission for your visit. Be aware that this permission
may not be granted.

Especially when using older guidebooks in this
collection, note that locations may have changed
drastically. Likewise, geological interpretations may
differ from current understandings.

Please respect any trip stops designated as "no
hammers", "no collecting" or the like.

Consider possible hazards and use appropriate caution
and safety equipment.

NEIGC and the hosts of these online guidebooks are not
responsible for the use or misuse of the guidebooks.

NEW ENGLAND INTERCOLLEGIATE GEOLOGICAL CONFERENCE

Informally organized in 1901 by William M. Davis, the
NEIGC provides an opportunity for students, faculty, and
industrial geologists to share the results of recent geologic
studies in New England. With the exception of eight years
during the two World Wars, the conference has met every fall
since the original trip to the Connecticut Valley led by
Prof. Davis. The 1980 meeting in Presque Isle is the 72nd
one held by the conference and the tenth meeting in Maine.
It was not until 1925 (in Waterville) that the conference
held its first meeting in Maine and the state was revisited
about every ten years until 1960 (1934, Lewiston; 1950,
Bangor; 1960, Rumford) . Since 1960 the conference has been
returning to Maine more frequently which attests to the
increasing amount of new work being done in New England's
largest state and the greater ease with which people can
travel to and within the state (1965, Brunswick; 1966,
Katahdin; 1970, Rangeley; 1974, Orono; 1978, Calais).
This year's meeting is the first in Aroostook County,
Maine's largest, and it is fitting that trips in neighboring
New Brunswick are included since the northwestern part of
the province has many geological and cultural ties to "The
County."

li

CONFERENCE ORGANIZATION

General Chairman: David C. Roy

Registration: Richard S. Naylor, Northeastern University

Local Arrangements: William H. Forbes, University of Maine,

Presque Isle
Guidebook Editors: David C. Roy and Richard S. Naylor

Trip Leaders

Ayuso, Robert A., Virginia Polytechnic Institute and State

University
Borns, Harold W. , University of Maine, Orono
Brewer, Thomas, Boston State College
Caldwell, D. W. , Boston University

Forbes, William H. , University of Maine, Presque Isle
Gauthier, Claude, Geological Survey of Canada
Genes, Andrew N. , Boston State College
Hamilton-Smith, Terence, Sohio Petroleum
Hon, Rudolf, Boston College
Rasper, Andrew E., Rutgers University
Kite, J. Steven, James Madison University

Lutes, Glen G. , Department of Natural Resources, New Brunswick
Naylor, Richard S., Northeastern University
Neuman, Robert B. , U. S. Geological Survey
Newman, William A., Northeastern University
Osberg, Philip H. , University of Maine, Orono
Pickerill, R. K. , University of New Brunswick
Rankin, Douglas W. , U. S. Geological Survey
Rast, N., University of Kentucky
Roy, David C. , Boston College

St. Peter, C, Department of Natural Resources, New Brunswick
Thibault, Jacques, Department of Natural Resources, New

Brunswick
Wones, David R. , Virginia Polytechnic Institute and State

University

Acknowledgment

The editors wish to thank Marjorie H. Roy for her
assistance in the final assembly of this Guidebook.

111

IN DEDICATION

Ely Mencher
1913 - 1978

This guidebook is dedicated to Ely Mencher
who was a friend, teacher, and colleague
to over a generation of students at MIT
and later at City College. His work from
1962 to 1976 and those of his undergraduate
and graduate students in northern Maine
established the foundation upon which much
of this conference is based. His quiet wit,
generosity, and sense of responsibility
endeared him to those of us who had the
privilege to meet and work with him.

IV

CONTENTS

Page

Conference Organization iii

Dedication to Ely Mencher iv

Contents v

Tectonics and Sedimentation in Northeastern Maine and
Adjacent New Brunswick, David C. Roy 1

Outline of the Pleistocene Geology of Northern Maine,

Andrew N . Genes 2 2

TRIP A-l : Geology of the Bottle Lake Complex, Maine,

Robert A. Ayuso and David R. Wones 32

TRIPS A-2 and B-l: Geology and Petrology of Igneous

Bodies within the Katahdin Pluton, Rudolf Hon 65

TRIPS A-3 and B-2: Alpine Glaciation of Mt. Katahdin,

D. W. Caldwell 80

Bedrock Geology of the Shin Pond-Traveler Mountain

Region, Robert B. Neuman and Douglas W. Rankin 86

TRIP B-3: The Traveler Rhyolite and Its Devonian

Setting, Traveler Mountain Area, Maine, Douglas W. Rankin 98

TRIP B-4: The Core of the Weeksboro-Lunfcsoos Lake

Anticline, and the Ordovician, Silurian, and Devonian

Rocks on its Northwest Flank, Robert B. Neuman 114

TRIP B-5: Plant Fossils (Psilophytes) from the Devonian
Trout Valley Formation of Baxter State Park, Andrew E.
Kasper, Jr 127

TRIP B-6: Ordovician and Silurian Stratigraphy of the
Ashland Synclinorium and Adjacent Terraine, David C. Roy 142

TRIP B-7 and C-2 : Bedrock Geology of the Presque Isle

Area, Richard S. Naylor 169

TRIP B-8: Wisconsinan Glaciation of Northern Aroostook
County, Maine, Andrew N. Genes and William A. Neuman 179

TRIP B-9: Deglaciation of the Edmundston Area and

Reappraisal of Glacial Lake Madawaska Interpretation,

Claude Gauthier and Jacques Thibault 188

v

TRIP B-10: The Geology and Deformation History of the

Southern Part of the Matapedia Zone and Its Relationship

to the Miramichi Zone and Canterbury Basin, N. Rast,

G. G. Lutes, and C. St. Peter 191

TRIP C-l: Stratigraphy and Sedimentology of the Siegas

Formation (Early Llandovery) of Northwestern New

Brunswick, Terence Hamilton-Smith 202

TRIP C-3: Biostratigraphic Trip Across Northeastern

Maine, William H. Forbes 212

TRIP C-4: Late-Glacial and Holocene Geology of the

Middle St. John River Valley, J. Steven Kite and Harold

W. Borns, Jr 213

TRIP C-5: Sedimentology of Silurian Flysch, Ashland
Synclinorium, Maine, David C. Roy 232

TRIP C-6: Wisconsinan Glaciation of Eastern Aroostook

County, Maine, Andrew N. Genes, William A. Newman,

Thomas B . Brewer 253

TRIP C-7: Structure and Sedimentology of Siluro-

Devonian Between Edmunston and Grand Falls, New

Brunswick, P. Stringer and R. K. Pickerill 262

TRIP D-l: Stratigraphic and Structural Relations in

the Turbidite Sequence of South-central Maine, Philip

H. Osberg 278

VI

TECTONICS AND SEDIMENTATION IN NORTHEASTERN
MAINE AND ADJACENT NEW BRUNSWICK

by

David C. Roy

Department of Geology and Geophysics

Boston College

The field trips of the 1980 NEIGC will examine the bedrock and surficial
geology of a large region in northeastern Maine and neighboring New Brunswick
as shown in Figure 1. The history of intensive geologic studies in this region
is more recent than for more southerly parts of New England but the results of
the studies are important to the "big picture" because the major rock belts
within the central part of the northern Appalachian range pass through it and
are seen at lower metamorphic grade. In fact, the best fossil-dated strati-
graphic cross-section of the range from the Connecticut Valley-Gaspe
Synclinorium through the Bronson Hill-Boundary Mountain Anticlinorium and across
the western flank of the Merrimack Synclinorium is present in northern Maine
and western New Brunswick.

The bedrock trips during the conference will not show the complete section
but will give participants a clear view of the complex stratigraphy from the
Bronson Hill-Boundary Mountain terrain across the Aroostook-Matapedia Belt, to
the Miramachi Belt of northern New Brunswick. The essential features of the
stratigraphy of the region are summarized in Figure 2. As will be illustrated
below, the early Paleozoic stratigraphy of the region of this year's conference
is similar to that seen at higher metamorphic grade during the 1970 Rangeley
conference and more recently during the 19 74 conference in Orono. The purpose
of this article is to provide an overview of the regional stratigraphy in the
region covered by the conference. In addition an attempt is also made to show
both the interplay of tectonics and sedimentation suggested by the stratigraphic
variations and to illustrate the broad regional continuity of the resulting
tectono-stratigraphic picture. It is hoped that the treatment will help partici-
pants to keep track of the formational "players" in the game and to begin to see
the positions they may have played during the evolution of this part of the
northern Appalachians.

Pre-Middle (pre-Caradocian) Ordovician Stratigraphy

In northern Maine our view of the pre-Middle Ordovician is very restricted.
A sparsely fossiliferous but relatively thick pre-Caradocian section has been
described by Neuman (1967; Neuman and Rankin, this volume) in the core of the
Weeksboro-Lunksoos Lake Anticline. These rocks include the Grand Pitch Form-
ation (varigated slate and quartzite) and the overlying Shin Brook Formation
(tuf faceous volcanic rocks and minor sedimentary rocks) . Rocks similar to the
Late Precambrian/Early Cambrian Grand Pitch have been mapped to the northwest
of the Weeksboro-Lunksoos Lake anticline by Hall (1970) as the Chase Brook
Formation. The lower quartzite-rich unit of the Tetagouche Group of the
Miramachi Anticlinorium in New Brunswick is commonly considered to also be
equivalent to the Grand Pitch (Rast, St. Peter, and Lutes, this volume).

c

u

0)

•u

co

co

W

*v

T3

C

CO

.H

M

C

W

S

CU

a

c

•H

CO

C

cd

•H

JS

O

crj

iH

crj

&.

CX

<

C

}-i

cu

,£

4-J

M

o

3

0)

.C

■u

>-H

o

CO

cu

•

U

M

d

o

4J

•H

ctt

IS

cu

CO

"4-1

d

3

o

>-(

•H

pa

c

o

5

4J

cu

o

2

cu

4-1

TJ

fi

■H

CO

CO

a.

#t

•H

CJ

o

CU

c

J3

•H

CU

M

0

P-l

Cf

cu
$-1

•H

I

AGE

(WEST)

FORMATIONS

(EAST)

Figure v2: Stratigraphy of northeastern Maine and adjacent parts of

New Brunswick. Star indicates that the unit is dated
by fossils.

The Miramachi belt is, however, separated from the Weeksboro-Lunksoos Lake
Anticline by younger rocks of the Aroostook-Matapedia Belt and it is therefore
difficult to establish lateral continuity of the quartzite-rich sequences.

The Shin Brook Formation overlies the Grand Pitch unconformably and
contains Late Early or Early Middle brachipods and trilobites (Neuman, 1967).
No clear lithologic correlatives of the Shin Brook have been described in Maine
or New Brunswick although largely slate-graywacke sections elsewhere are probable
temporal equivalents (Neuman, 1967; Hall, 1969). The slate-siltstone-graywacke
unit of the Tetagouche Group (Rast, St. Peter, and Lutes, this volume) may be
in part correlative with the Shin Pond.

The unconformity at the base of the Shin Pond Formation may represent a
major hiatus and has been attributed to the Penobscott Disturbance by Neuman
and Rankin (1966). An unconformity in apparently similar stratigraphic position
may be widespread in the northern Appalachians (Hall, 1969), but good dating of
the unconformity is not everywhere possible and the deformation and erosion
suggested by it may not represent a widespread synchronous event.

The paleogeography of the earliest Paleozoic is not well understood as
might be expected from the limited distribution of these poorly dated rocks of
pre-Middle Ordovician age. Since these rocks are found in anticlinorial tracts
that are separated by synclinorial belts of younger rocks it is difficult to
access regional patterns. Reliance on lithologic correlations, especially across
the regional structure, is dangerous because rapid cross-strike facies changes
are common in younger rocks (Roy and Mencher, 1976) and are likely in the
pre-Middle Ordovician as well. It seems clear that the Grand Pitch-type
lithofacies in which abundant quartzose sandstone is interlayered with slate
(in variable proportions) is widespread in the northern Appalachians and
probably underlies much of the region of this conference. The abundance of
quartzite in the lithofacies suggests sialic (cratonal?) sources along the
margins of and possibly within the orogen. The variations in abundance of
quartzite may reflect changes in proximity to the sialic sources.

Late Ordovician (Caradocian/Ashgillian) Lithofacies

Beginning with rocks of the later part of the Ordovician it becomes
possible to work out fairly clear lithofacies patterns in northeastern Maine and
parts of New Brunswick. Such a lithofacies pattern involving three major litho-
facies and representing the Caradocian and Ashgillian is shown in Figure 3.1

1In figures 3,4,5, and 6 the following information applies. Towns and cities
are abbreviated as follows: A, Ashland; Aug, Augusta; B, Bathurst; Ban, Bangor;
C, Caribou; Cal, Calais; D, Dalhousie; F, Fredericton; FK, Fort Kent; G, Green-
ville; GF, Grand Falls; H, Houlton; MT. K. , Mount Katahdin; M, Matapedia;
P, Portage; PI, Presque Isle; R, Rockland; VB, Van Buren. Primary sources of
information are: Roy, 1970; Roy and Mencher, 1976; Pavlides, 1965, 1968, 1971,
1972, 1973; and Hall, 1970; Neuman, 1967; Boudette and others, 1976; St. Julien
and Hubert, 1975; Boudette and Boone, 1976; Boone, 1973; Anderson, 1968; Hamilton-
Smith, 1969, 1970; Boucot, 1961, 1969a, b; Boucot and others, 1964, Moench, 1970a,
b, 1971, 1973; Ludman, 1976; (foot-note continued on the next page)

z

•
•

•
E

<

E

JC

e

o

»'

o

>

-

Q CO

s yj

. o

oo

1°

yj if

M

o-!-o

cr-1

/

O

/
/

-D

/

< >

/

>

/

\

\

3

o

M
O

k

I

J

i

V

a
»

I i

r
<

E
■

a.

s

3

m
3

u

<

o

m

I 9 ° * I •
5 5 2 * X

> CO » > X ™

E3DE3EE1 [

o
a

■
a

II

I + 'ff^S , l\ *

ro /' .1' I Y.

to

h

*5

> > * > , -I 4/ 'A

ft fi*$ <,V

in t » • » *. •

O V . .: :■:■:.:% \.i 5

S

0)

C
•H

S

CO

cu

•H

o
cd

O

>
> >

•I
oo

The volcanic lithofacies forms a broad tract in northwestern Maine.
The western limit of the lithofacies in much of northwestern Maine is obscured
by Siluro-Devonian rocks but it is possible that the lithofacies is continuous
in the subsurface with the Ascot-Weedon volcanic sequence in Quebec. In
northern-most Maine the volcanic lithofacies is represented by the Winterville
Formation (Roy, Trip B-6» this volume) but the lithofacies is embodied in a
variety of named and un-named units to the southwest as indicated in Figure 3.
The volcanic lithofacies in Maine consists largely of fragmental and pillowed
spilitic basalts, dacite, karatophyre, soda rhyolite and apparently less
abundant interstratif ied black sulfidic slate, chert, and graywacke. Hynes
(19 76) assigns the volcanic rocks of the Winterville and Bluffer Pond forma-
tions to the alkali olivine basalt suite based on titanium concentrations in
pyroxene. Hynes has found the un-named volcanic rocks of the Weeksboro-Lunksoos
Lake Anticline to be more varied in silica content and to be transitional
between alkaline and non-alkaline rocks.

Variably altered Mafic and ultramafic rocks are found in the volcanic
belt southwest of the region of the conference. As discussed by Boudette and
Boone (1976) some of these rocks, especially those marginal to the Chain Lakes
Massif, may be part of a dismembered pre-Middle Ordovician ophiolite sequence;
others appear to be younger intrusions. Massive sulfide deposits of the Kuroko
type are present locally within the volcanic lithofacies but have not been
described in detail as yet.

East of the Volcanic Lithofacies is a belt of more or less contemporaneous
slate and graywacke. The slate-graywacke lithofacies is found in the Madawaska
Lake Formation in the north (Figure 3) but is also seen in the Chandler Ridge
(Pavlides, 1968) and Mattawamkeag (Ekren and Frischknecht , 1967) formations
farther south. The Mars Hill Conglomerate (Pavlides, 1978) may be a part of this
lithofacies located beneath the limestone-rich Carvs Mills Formation of the
Aroostook-Matapedia Belt. The Quimby-Greenvale Cove sequence of western Maine
is considered here to be expressions of this lithofacies. The graywacke beds
of this belt contain abundant volcanic detritus which presumeably is derived
from volcanic islands in the western volcanic belt. In the vicinity of Portage,
Maine, the Winterville Formation interfingerTs with the Madawaska Lake Forma-
tion as discussed by Roy (Trip B-6, this volume). No evidence of penecomtemp-
oraneous deformation (e.g. subduction zone tectonism) is evident along the
lithofacies boundary at Portage or elsewhere in Maine.

East of the slate-graywacke lithofacies and forming much of the anticlin-
orial terrain of the Aroostook-Matapedia belt is the thick limestone-rich Carys
Mills Formation. The lower part of this formation is known to represent the
Upper Ordovician (Pavlides, 1968) and thus it forms the next lithofacies. In the
Presque Isle-Caribou area Roy (1978; Trip B-6, this volume) has found the lower
part of the Carys Mills to be more graywacke- rich than the upper part. Both the
slate and graywacke of this lower phase are calcareous and are interlayered with

Ludman and Griffin, 1974; Pankiwskyj and others, 1976; Osberg, 1968; Gates,
1961; Ruitenberg and Ludman, 19 78; Noble, 1976; Lee and Noble, 1977; Greiner,
1973; Greiner and Potter, 1966; Rast and Stringer, 1974; Ayrton, and others,
1969 i Skinner, 1964; Naylor and Boucot, 1965; Pavlides and others, 1968; Lajoie
and others, 1968.

minor micritic limestone. The lower Carys Mills forms the slate-graywacke-
limestone lithofacies of Figure 3. The lithofacies may extent into central
Maine but it can only be inferred to extend a short distance south of Houlton
based on the distribution of the Carys Mills Formation as a i^hole.

East of the Carys Mills belt, largely in New Brunswick, rocks of Late
Ordovician age are generally assigned to the Tetagouche Group. Most of the
subdivisions of the Tetagouche Group as originally described by Helmstaedt (1971)
for the Bathurst-Newcastle area have been extended into southwestern New
Brunswick by Venugopal (19 79) and Rast, Lutes, and St. Peter (this volume). In
the Miramichi zone near Woodstock, N.B., slate and graywacke units of the
Tetagouche Group appear to comprise the section above the quartzite-rich
sequence that forms the basal unit of the group. In so far as these slate-
graywacke units are Late Ordovician in age, they may represent an analogous
lithofacies to that found west of the Aroostook-Matapedia belt. Presumeably
the eastern slate-graywacke lithofacies interfingers with the mafic volcanics
of the "type" Tetagouche in the Bathurst area. At present, however, it is
difficult to be certain of lithofacies relationships east of the Aroostook-
Matapedia belt because of major faults along the northwestern margin of the
Miramichi Anticlinorium (Rast and Stringer, 1974) and the paucity of fossils
within the anticlinorium.

The Taconian Orogeny

The Taconian Orogeny in northeastern Maine was a relatively mild deform-
ational event but it produced substantial changes in paleogeography and
sedimentation in the Aroostook-Matapedia basin. The region of the conference
is southeast of the Thetford Ultramafic trend (Figure 3) that marks the suture
zone of the Taconian Orogeny as invisioned by St. Julien and Hubert (1975).
The southeastern extent of Taconian deformation in Maine appears to coincide
roughly with the transition from the volcanic to the slate-graywacke lithofacies
In the eastern townships of Quebec the Taconian began in the Early Ordovician
and proceeded through the Late Ordovician with the development of large-scale
overthrust tectoncis apparently associated with a complicated obduction/subduc-
tion history. In northern Maine the Taconian was less dramatic and is seen to
be younger, namely latest Ordovician and earliest Silurian (Roy, Trip B-6, this
volume) .

No metamorphic fabric of Taconian origin has been reported in northern
Maine from any of the lithofacies belts. Cleavage in Ordovician slate is
essentially axial planar to Acadian folds. Volcanic rocks and graywacke beds
of Ordovician age (and younger for that matter) are generally incleaved north
of the lattitude of Presque Isle but show increasing foliation along strike to
the south. The cleavage in these more competent rocks is also consistent with
Acadian deformation.

Richter and Roy (1974, 1976) argue that Taconian metamorphism of the
Winterville Formation probably did not exceed prehnite-pumpellyite grade.
Assessments of Taconian metamorphism elsewhere in northern Maine have not been
made and become more difficult as Acadian metamorphic grade increases southwest-
ward.

The rocks of the Aroostook-Matapedia basin have produced no clear evidence

8

of Taconian deformation. Early recumbent folds reported by Rast, Lutes, and
St. Peter (this volume) in the Carys Mills near Woodstock represent pre-Acadian
deformation that may be as old as Taconian. However, it is also possible that
these early folds are non-tectonic and represent syndepositional slump folds as
described elsewhere in the Carys Mills (see Stringer and Pickerill, this volume)
or they may be tectonic folds associated with Salinic deformation at the close
of the Silurian (see below) .

Taconian deformation of the Miramichi zone is well established (Rast and
Stringer, 1974; Rast and others, this volume) and probably carried the Tetagouche
rocks to at least the greenschist grade locally (Helmstaedt, 1971). Skinner (19 74)
has proposed an early northwest-trending fold system followed by the generation

of younger more open folds with northeast-trending axes for the Bathrust area but
he does not attempt to assign ages to the fold events.

Lower Silurian Lithofacies

Following the Taconian Orogeny substantial paleogeographic changes are
revealed by Llandoverian lithofacies. Figure 4 shows the distributions of the
major lithofacies for the Late Llandoverian-Early Wenlockian time period and the
approximate limits of emergent areas. A paleogeographic map for the Early
Llandoverian produced by Ayrton and others (1969) is similar to that shown for
a slightly later time period in Figure 4. One major difference between the early
and late Llandoverian lithofacies pictures is the presence of the upper "ribbon-
rock" facies of the Carys Mills in the interior of the Aroostook-Matapedia basin
during the earlier part of the Llandovery. The Late Llandoverian-Wenlockian
interval was selected for Figure 4 because of the larger amount of paleontologic
data available and greater complexity of the lithofacies pattern.

The land area called "Taconia" was probably contiguous with the pre-
Taconian North American Craton since the ocean-basin closure described by St.
Julien and Hubert (19 75) appears to have completely removed marine conditions
from northwestern Maine and adjacent Quebec. The thick sequence of coarse
clastic sediments of the Frenchville-Rangeley belt represents deposits along the
newly established continental margin in Maine. In northeastern Maine these
coarse clastic sequences give way eastward to much finer grained deep-water
lithofacies in the Aroostook-Matapedia Basin (Roy, Trip B-6 and C-5 , this volume).
Similar off-shore lithofacies changes can be seen from western to central Maine
along the western flank of the Merrimack Synclinorium.

The Miramichi terrain appears to have also been uplifted by Taconian de-
formation as evidenced by coarse sediments deposited in the Bathurst Basin
(Noble, 1976). The extent of "Miramichia" in southwestern New Brunswick is not
well known and faulting has been largely responsible for limiting our understand-
ing of the detail lithofacies picture in the Woodstock area.

It is of interest to note that by Late Llandoverian time the broad Merrimack
Synclinorium in Central Maine seems to have been divided by Miramichia into two
basins, the Aroostook-Matapedia Basin and the Fredericton Basin, in New Brunswick
as pointed out by McKerrow and Ziegler (1971) who figured a map similar to that
shown in Figure 4. McKerrow and Ziegler (1971) suggest that the Fredericton
Basin was the main oceanic continuation of the Merrimack Basin and that the
Aroostook-Matapedia Basin was a branch that terminated in the Gaspe Peninsula.

10

As shown in Figure 4, the relationships of the deep-water facies of the
Aroostook-Matapedia Basin in Maine to more shallow-water sedimentary sequences
in the Bay of Chaleur are unclear. The northeastern Maine basinal Silurian is
similar to the sequence in the Temisconta-Matapedia region of Quebec as described
by Lajoie and others (1968) that are here shown as deposited in the "Mistigoug-
ueche Basin". The Silurian of the Mistigougueche Basin was probably everywhere
deposited on Taconian-deformed pre-Silurian rocks unlike the basinal sequences
in Maine. It is possible that the Silurian Aroostook-Matapedia sequence in
Maine is continuous with the Bathurst Basin as described by Noble (1976).

Late Silurian Lithofacies

By the Late Wenlockian as shown in Figure 5, elevation of the Taconian
land areas had been greatly reduced by erosion and possibly by subsidence.
Everywhere in the deep-water basins that were established in the Early Silurian
the deposits of Late Silurian age indicate more distal environments. In addi-
tion, shallow-water deposits appear to be widespread in the previously up-
land areas, especially during the Ludlovian.

Small islands, including Somerset Island of Boucot (1961, 1969), were
probably common on the inundated Taconia as Late Silurian transgression took
place. These islands provided the coarse detritus for sandstones and conglomer-
ates that are common in the shallow-water sequences. Intermediate and mafic
volcanic rocks interlayered with the shallow-water sediments are also common and
suggest at least some tectonism during the advance of marine conditions.

Eastward in the basin one finds a considerable reduction in conglomerate
deposition and the development of generally thin-bedded flysch typified by the
Jemtland, Upper Smyrna Mills, Upper Allsbury, and Upper Sangerville formations.
The sandstone beds within the basins are predominantly calcareous, rusty-
weathering quartzofeldspathic graywacke. The graywacke beds become increasingly
more lithic and abundant toward the west. The Late Silurian sandstones on the
west side of the basin, extending from Augusta to beyond Van Buren (Figure 5),
are clearly derived from the somewhat subdued Taconia to the west. Derivation
of similar sandstones in the southeastern part of the Merrimack Basin from an
eastern "platform" may have begun in the Early Silurian and continued into the
Ludlovian (Figure 5; McKerrow and Ziegler, 1971; Ludman and Griffin, 1974;
Ruitenberg and Ludman, 1978). The two-sided character of the Aroostook-Matapedia
Basin is less well documented for the Late Silurian than for the Early Silurian
(Figure 4) but is probable. Derivation of Late Silurian detritus from Miramichia
in the Bay of Chaleur area is well known (Greiner and Potter, 1966; Greiner, 19 73;
Noble, 19 76; Lee and Noble, 1977) and Anderson (1968) reports clasts of Miramichia
slate and graywacke in Silurian conglomerates northeast of Woodstock, New Brunswick.

The Salinic Disturbance and Earliest Devonian Lithofacies

The Salinic Disturbance (Boucot, 1962) is indicated by a disconformity in
the Presque Isle area between the Jemtland Formation and Early Devonian (Late
Gedinnian) volcanic and sedimentary rocks of the Dockendorf Group (Boucot and
others, 1964). Unconformities between shallow-water Late Silurian (Ludlovian)
sedimentary rocks and quite similar rocks of Late Gedinnian age are common also
along the Munsungun-Pennington Mountain Anticlinorium (Hall, 19 70; Boone, 19 70;

11

cd

cd

c

cd

d
<u
o

cd

cd

d

cd

<u
d

•H

cd

d

•H

co

<U
•H
O

cd
o

cd

•H
>
O

iH

d

hJ

i

d

cd

•H

M
o
o

iH

d

cu

&

■M

cd
.-J

o

cd

6

<4-l

o

d
o

•H

4-1

d

•H
4->

w

•H

Q

in

50
•H

12

^_«^_

•H

To

i

•H

T3

p.

a)

•

o

|1

a)

3

4J

60

o

cd

Cd

§

u

4-1

c
a)
o
cd

cd

CD

d

•H

id

CD
(1)
•H
O

cd
o

X!

I

cd

o

cd

6

c

O
•H
4J

x>

•H
U

4-1

CO

•H
Q

CO

60
•H

0)

o

D

4-1
CO
•H
Q

O
•H

fl

•H

cd

C/3

CU

J3

4J

13

Roy and Mencher, 1976). Between these two regions where unconformities between
Silurian and Devonian rocks are present there is a belt in which Devonian rocks
apparently rest conformably on the Silurian. In the Ashland Synclinorium the
Fogelin Hill Formation which probably contains the systemic boundary rest con-
formably on the Jemtland (Roy and Mencher, 1976). Hall (1970) suggests that
the Siluro-Devonian boundary lies within the Third Lake Formation in the south-
eastern part of the Spider Lake region. It is therefore possible that an
elongate basin in northeastern Maine survived the Salinic uplifts which caused
erosion both to the northwest and southeast of the basin as shown in Figure 6. .

The extents of the syn-Salinic basin are not well established. To the
northeast the basin may correspond to the lower Cape Bon Ami Formation which
seems to rest conformably on the St. Leon Formation (Lajoie and others, 1968).
To the southwest it is not possible to trace the "axis" of the basin very pre-
cisely or even to be sure of its persistence beyond the Third Lake area. The
axis is presumed to have passed northwest of the position of Mount Katahdin
during Predolian-Early Gedinnian time. By Late Gedinnian (Figure 6) the basin
appears to have widened considerably. It is generally presumed that the slate-
rich sequence in central Maine extends up into the Devonian. By the Late
Gedinnian (and possibly earlier) a connection between the northern basin in
which the Fogelin Hill was deposited and a basin in central Maine is possible.
The dimensions of the central Maine basin are uncertain and depend on the true
extent of Early Devonian rocks in central and eastern Maine as discussed
below.

Much of central and southern New Brunswick was probably uplifted during
the Salinic interval (Figure 6) . This conclusion is based on the apparent
absence of sedimentary rocks of Early Devonian age in that region and the pre-
sumption that the "Kingsclear Series" of the Fredericton Basin does not extend
into the Devonian. The present writer believes that this emergent area extended
well into southeastern Maine because it is suspicioned that the Devonian is
absent there also. Boucot (1968) has essentially come to the same conclusion.
Unfossiliferous formations such as the Flume Ridge, Bucksport, Fall Brook, and
Brighton are lithologically more similar to dated Silurian units farther north
and northwest than they are with dated Devonian sections. In addition, the early
recumbent folds postulated by Osberg (this volume) may be Salinic folds in a
completely Silurian section that were later refolded by more upright Acadian
folds. If the units just mentioned are indeed restricted to the Silurian, the
syn-Salinic basin may have been restricted to west-central Maine where the Madrid,
Solon, and similar units are present.

whatever the extent of the eastern land area, an important belt of volcanic
and sedimentary rocks developed along northwestern edge in New Brunswick and
northeastern Maine. The Dalhousie Group of the Bay of Chaleur region and the
Dockendorf Group of the Presque Isle area lie along a more or less continuous
and broad volcanic belt that is immediately post-Salinic (Boucot and others,
1964) . Naylor (this volume) suggests that this belt represents a continental
volcanic arc. The Dockendorf Group has been shown to be about 3600 meters thick
(Boucot and others, 1964) and the Dalhousie section is at least 600 meters thick
and possibly thicker. Substantial subsidence must have accompanied the forma-
tion of these thick sections of volcanic rocks and interlayered shallow-water
sedimentary rocks.

Rapid westward facies changes are present in the Late Gedinnian rocks

14

between Presque Isle and Ashland. In Ashland mudstone conglomerate, limestone,
polymictic conglomerate and sandstone beds are interlayered with slate. The
mudstone conglomerates are usually monomicitc with clasts of either volcanic
rocks or clasts from the Jemtland Formation (some containing graptolites) .

In the Matagamon Lake area (see Neuman and Rankin, this volume) there
is good evidence of a large delta involving the Seboomook Formation and
Matagamon Sandstone (Hall and others, 1976). Hall and others (9176) conclude
that the deltaic sediments were derived from the east beginning in the
Gedinnian, and that the delta prograded westward during the Siegenian. The
delta implies a major upland to the east and the presence of a relatively deep
basin offshore to the northwest. Additional such deltas with coarsening-
upward sections may be present in western Maine (e.g. the Madrid-Carrabassett
sequence of Boone, 1973). Such delta sequences may reflect continuation of
the southeastern source area into the Maine Slate Belt.

Siegenian Time and The Seboomook Formation

The term "Seboomook" has been applied over the years to the widespread
gray slate and graywacke sequences in northwestern Maine. Boucot (1970) has
reviewed the distribution of such sequences in Maine and elsewhere within the
northern Appalachians and summarized the paleontological control on their ages.

In the Ashland-Portage area and in much of northwestern Maine, the
Devonian section is divisible into two parts. The lower part is a lithologically
variable section made up of conglomerate, lithic sandstone and limestone inter-
layered with cleaved mudstone or slate. This lower part usually is fossiliferous
(both fauna and flora) and up to about 1 km or so thick. The upper part consists
of generally fine-grained, well-cleaved slate with minor but typically cyclically
interlayered thin beds of graywacke. The upper part is unfortunately rarely
fossiliferous and has no defined "top".

The Seboomook of the Matagamon Lake Area (Hall and others, 1970) and the
Moose River Synclinorium (Boucot, 1961, 1969), on the other hand, coarsens
upward and is succeeded by fossiliferous shallow-water sandstone units (Mata-
gamon and Tarratine sandstones) followed by silicic volcanic sequences
(Traveller Rhyolite and Tomhegan Formation). These stratigraphically "topped"
Seboomook sections are difficult to reconcile with the Seboomook further north
and west that seems to not have a definable stratigraphic top and indeed may
have been "terminated" by the Acadian Orogeny. As pointed out by Boucot (1970),
very careful mapping within this monotonous sequence must be done in order to
subdivide the now broad "Seboomookland" into subbasins and sources of sediment
supply.

The work of Hall and others (1970) is just such an effort. At the time
of his death, Ely Mencher was also attacking this very problem in a broad sub-
division of "Seboomookland" west of Ashland. Ely's approach was to try and
piece together the facies variations in the lower, better-dated part of the
section in a large region where fossils were likely to be found due to low
regional grade. Ely's field style and meticulous attention to detail were well
suited to the task and it is a shame he could not finish it.

15

The Acadian Orogeny

The principal fold and cleavage producing event in northern Maine
post-dates Siegennian deposition and is assigned to the Acadian Orogeny. The
Mapleton Formation of Late Middle Devonian age (Schopf, 1964; White, 1975;
Boucot and others, 1964; see Naylor, this volume) rests unconformably on nearly
vertical lower Devonian and Silurian rocks and provides the best stratigraphic
evidence for the end of the orogenic phase of the Acadian.

Folds

The style of Acadian folding varies from place to place depending pri-
marily on the rocks involved in the folded sections (Pavlides and others, 1964;
Roy, 1970; Pavlides, 1974). Where thick volcanic or sandstone-conglomerate
sections are present the folds are more open and concentric and flexure-slip
mechanisms predominated. The Chapman Syncline (involving the Dockendorf Group) ,
the Stockholm Mountain Syncline (involving the Frenchville Formation) , and the
Castle Hill Anticline (involving both the Winterville and Frenchville formations)
are good examples of major folds that are simple in form (see Figures 1 and 2,
Roy, Trip B-6, this volume). Where the stratigraphic section is dominated by
pelite more tightly appressed, generally symmetric and steeply plunging folds
of either concentric or similar form are present. Pelite mobility in these
folds is indicated by common pelite injection along cleavage planes in limestone
and graywacke beds as well as hinge thickening of pelite beds (Pavlides, 1965;
Roy, 1970).

Disharmonic folding is seen locally (Pavlides, 1973) and is common on an
outcrop scale in the laminated Silurian ironstones (Roy, 1970). Disharmonic
folding on a large scale has been suggested for the Chapman Syncline (less
competent pre-Devonian versus the competent Devonian volcanic rocks) by Pavlides
(1974) and may be characteristic of the Pennington Mountain Anticlinorium where
a thick previously deformed and volcanic-rich Winterville Formation is overlain
by the thick slate-rich Seboomook Formation.

Pavlides (1965, 1971, 1972) has documented complicated fold geometries
(including overturned fold plunges) in the Ordovician and Silurian pelite-rich
units in the Aroostook-Matapedia Anticlinorium. Outcrop size and spacing have
so far precluded complete analysis of possible pre-Acadian folding in these
rocks, but they are generally thought to have undergone only one deformation
(Pavlides, 19 74, Roy, 1970). Early folds may be more common than previously
recognized, but reports are rare (Hamilton-Smith, 1970; Rast, Lutes and St.
Peter, this volume) and their distinction from non-tectonic folds is difficult.

Cleavage

A prominant S cleavage is present in the pelitic rocks. The S cleavage

is generally parallel or nearly parallel to axial planes of Acadian folds.
Massive sandstone and conglomerate beds or volcanic rocks in the northern part
of the conference area typically show no cleavage; thin graywacke beds in slate-
rich sections, however, do show a fracture cleavage. Southward from the
lattitude of Presque Isle sandstone beds and volcanic rocks show increasingly
well developed foliation.

16

Cleavage in pelitic intervals varies from fracture cleavage in which there
is a low degree of micaceous alignment to flow cleavage in which there is perva-
sive orientation of micaceous minerals parallel to cleavage surfaces. In
subgreenschist slates fracture cleavage is characteristic of non-calcareous
slate whereas flow cleavage is typical of calcareous slate. Calcareous slates
are usually phyllitic in appearance. Pressure solution effects, as discussed by
Stringer and Pickerill (this volume), are probably of widespread importance in
cleavage formation, especially in low-grade slates.

A late northwest trending fracture cleavage (S~) is associated with the
broad Houlton Oroflex which folds the axial surfaces of the Acadian folds in
the vicinity of Houlton (Pavlides, 1974).

Metamorphism

Acadian metamorphism in the northern part of the region of the conference
did not exceed prehnite-pumpellyite grade (Combs and others, 19 70; Richter and
Roy, 1976). The grade there increases from the prehnite-analcime subfacies in
the northwest (Pennington Mountain Anticlinorium) to the pumpellite-epidote-
Actinolite subfacies to the south and southeast. Metamorphism continuous to
increase southward from Presque Isle so that in the Bridgewater-Houlton area
low greenschist grade is present (Pavlides, 1965, 1971). Similarly, to the
south and southwest, grade becomes low greenschist in the Spider Lake area
Hall, 1970) and in the Weeksboro-Lunksoos Lake Anticline (Neuman, 1967).

The prehnite-pumpellyite paragenesis in the north, which is seen only in
volcanic rocks and in lithic graywacke beds, appears to be prograde. It is not
clearly associable with an Acadian fabric since the rocks containing the diag-
nostic minerals are not foliated. A regional increase in fluid C0? content
appears responsible for suppression of the diagnostic minerals of the facies
in volcanic rocks of suitable composition in the Presque Isle area (Richter and
Roy, 1976). Confirming the southeastward temperature increases suggested by
the changes in diagnostic sub-greenschist minerals are preliminary paleotempe-
rature results by Anita Harris of the USGS (Personal communication, 19 76) on
Silurian and Devonian conodonts. Her results suggest a temperature increase
from 50°C (lower Devonian Square Lake Limestone) in the prehnite-analcime zone
to 190-250°C within the pumpellyite-epidote-actinolite zone in the Castle Hill
Anticline and Chapman Syncline (Lower Silurian Frenchville and Spragueville
formations) .

Plutonism

Widely scattered granitic Acadian plutons are present in the area covered
by the conference. The two largest plutons are the Katahdin Quartz Monzonite
(Hon, this volume; Neuman and Rankin, this volume) and the Bottle Lake Complex
(Ayuso and Wones, this volume); both will be examined during the conference.
Worth of these large plutons are several smaller intrusives: the Nickerson Lake
and Pleasant Lake plutons south of Houlton (Pavlides, 1971); the Munson Pluton
just northwest of Presque Isle (Boucot and others, 1964); the Deboulli Stock
northwest of Portage (Boone, 1962); and the Chandler Lake Pluton southwest of
Ashland.

All of the plutons cut the Acadian cleavage and have well-defined met amor-
phic aureoles. For the most part the plutonic rocks are not strongly foliated.

17

Acknowledgements

Early versions of some of the material presented in this paper were
reviewed by J. Rehmer, J.C. Hepburn, C.J. Roy, and R.S. Naylor who gave
valuable suggestions. Sally Sargent gave able assistance in the preparation
of Figures 2,3,4, and 5.

References

Anderson, F.D., 1968, Woodstock, Millville and Coldstream map areas, Carleton
and York Counties, New Brunswick: Geological Survey of Canada, Mem. 353.

Ayrton, W.G.; Berry, W.B.N. ; Boucot, A.J.; Lajoie, Jean; Lesperance, P.J.;

Pavlides, Louis; Skidmore, W.B., 1969, Lower Llandovery of the northern
Appalachians and adjacent regions: Geol. Soc. America Bull., v. 80,
pp. 459-484.

Boone, G.M. , 1962, Potassic feldspar enrichment in magma: origin of syenite in
in Deboullie district, northern Maine: Geol. Soc. America Bull. , v. 73,
pp. 1451-1476.

1970, The Fish River Lake Formation and its environments of deposition;

in Shorter contributions to Maine geology: Maine Geol. Survey Bull. 2 3,
pp. 27-41.

19 73, Metamorphic stratigraphy, petrology, and structural geology of the

Little Bigelow Mountain Map Area, Western Maine: Maine Geol. Survey Bull.
24, 136p.

Boucot, A.J., 1961, Stratigraphy of the Moose River Synclinorium, Maine: U.S.
Geol. Survey Bull. 1111-E, 188p.

1962, Appalachian Siluro-Devonian: in Coe, Kenneth, ed., Some Aspects of

the Variscan Fold belt, 9th Inter-Univ. Geol. Congr. , Manchester Univ.
Press, pp. 155-163.

1968, Silurian and Devonian of the northern Appalachians, in Zen, E-an;

White, W.S.; Hadley, J.B.; and Thompson, J.B., Jr., eds. , Studies of
Appalachian geology; northern and maritime: New York, Interscience Pubs.,
Inc., chap. 6, pp. 83-94.

1969a, Silurian-Devonian of Northern Appalachians-Newfoundland: in_ Kay,

Marshal, ed. , North Atlanitc-geology and continental drift; Amer. Assoc.
Petroleum Geol. Mem. 12, pp. 477-483.

1969b, with contributions by Heath, E.W., Geology of the Moose River and

Roach River Synclinoria, northwestern Maine: Maine Geol. Survey Bull. 21,
117p.

, Field, M.T.; Fletcher, R. ; Forbes, W.H. ; Naylor, R.S.; and Pavlides, L. ,

1964, Reconnaissance bedrock geology of the Presque Isle quadrangle, Maine:
Maine Geol. Survey Quad. Mapping Serv. , no. 2, 123p.

18

Boudette, E.L. and Boone, G.M. , 1976, Pre-Silurian stratigraphic succession in
central western Maine: in Page, L.R., ed., Contributions to the strati-
graphy of New England: Geol. Soc. Amer. Mem. 148, pp. 79-96.

, Hatch, N.L., Jr. and Harwood, D.S., 1976, Reconnaissance geology of the

upper St. John and Allagash River basins, Maine: U.S. Geol. Survey Bull.
1406, 37p.

Coombs, D.S.; Horodyski, R.J.; and Naylor, R.S., 19 70, Occurrence of prehnite-
pumpellyite facies metamorphism in northern Maine: Amer. Jour. Sc. ,
v.368, pp. 142-156.

Ekren, E.B. and Frischknecht , F.C., 1967, Geological-geophysical investigations
of bedrock in the Island Falls quadrangle, Aroostook and Penobscot
counties, Maine: U.S. Geol. Survey Prof. Paper 527, 36p.

Gates, Olcott, 1961, The Geology of the Cutler and Moose River Quadrangles,

Washington County, Maine: Maine Geol. Survey, Quad. Mapping Series, No. 1,
67 p.

Greiner, H.R., 1973, Pointe Verte to Tide Head, Caleur Bay Area, New Brunswick:
in Rast, N. , ed. , Geology of New Brunswick, 1973 New England Intercolle-
giate Geol. Conf. Guidebook, pp. 58-70.

and Potter, R.R., 1966, Silurian and Devonian stratigraphy, Northern New

Brunswick: in Poole, W.H., ed. , Geology of parts of Atlantic Provinces,
Geol. Assoc. Can. Guidebook, pp. 19-32.

Hall, B.A., 1969, Pre-Middle Ordovidian unconformity in northern New England

and Quebec: in Kay, Marshall, ed. , North At Ian tic- geology and continental
drift: Amer. Assoc. Petroleum Geol. Mem. 12, pp. 467-476.

1970, Stratigraphy of the southern end of the Munsungun anticlinorium,

Maine: Maine Geol. Survey Bull. 22, 63p.

, Pollock, S.G. and Dolan, K.M. , 1976, Lower Devonian Seboomook Formation

and Matagamon Sandstone, Northern Maine: A Flysch Basin-Margin Delta
Complex: in Page, L.R., ed. , Contributions to the stratigraphy of New
England: Geol. Soc. America Mem. 148, pp. 57-63.

Hamilton-Smith, Terence, 1969, Paleogeography of northwestern New Brunswick

during the Llandovery: a study of the Provenance of the Siegas Formation:
Can. Jour. Earth Sci. , v. 8, pp. 196-203.

1970, Stratigraphy and structure of Ordovician and Silurian rocks of the

Siegas area, New Brunswick: New Brunswick Dept. Nat. Resources Mineral
Resources Br. Rept. Inv. 12, 55p.

Helms taedt, Herwart, 1971, Structural Geology of Portage Lakes Area, Bathurst-
Newcastle District, New Brunswick; Geol. Survey Canada Paper 70-28, 52p.

Hynes, Andrew, 1976, Magmatic affinity of Ordovician Volcanic rocks in northern
Maine, and their tectonic significance: Amer. Jour. Sci., v. 276, pp. 1208-
1224.

19

Lajoie, J.; Lesperance, P.J.; and Beland, J., 1968, Silurian stratigraphy and
paleogeolgraphy of Matapedia-Temisconata region, Quebec: Amer. Assoc.
Petroleum Geol. Bull., v. 52, pp. 615-640.

Lee, H.J. and Noble, J. P. A., 1977, Silurian stratigraphy and depositional

environments: Charlo-Upsalquitch Forks area, New Brunswick: Can. Jour.
Earth Sci. , v. 14, pp. 2533-2542.

Ludman, Allan, 19 76, Fossil-based stratigraphy in the Merrimack Synclinorium,
Central Maine: in Page, L.R. , ed. , Contributions to the stratigraphy of
New England: Geol. Soc. Amer. Mem. 148, pp. 65-78.

, Griffin, J.R., 1974, Stratigraphy and structure of central Maine: in

Osberg, P.H., Guidebook for field Trips in east-central and north-central
Maine: 1974 New England Intercollegiate Geol. Conf. Guidebook, pp. 154-179.

and Ziegler, A.M., 1971, The Lower Silurian paleogeography of New Bruns-
wick and adjacent areas: Jour. Geol., v. 79, pp. 635-646.

Moench, R.H. , 1970a, Evidence for premetamorphic faulting in the Rangeley

quadrangle, western Maine: jLn Boone, G.M., ed. , Guidebook for field trips
in the Rangeley Lakes-Dead River Basin region, western Maine: 1970 New
England Intercollegiate Geol. Conf. Guidebook, pp. D1-D12.

1970b, Premetamorphic down-to-basin faulting, folding and tectonic de-
watering, Rangeley area, western Maine: Geol. Soc. America Bull., v. 81,
pp. 1463-1496.

1971, Geologic map of the Rangeley and Phillips quadrangles, Franklin and

Oxford counties, Maine: U.S. Geol. Survey Misc. Geol. Inv. Map 1-605.

1973, Down-Basin Fault-fold tectonics in western Maine, with comparisons

to the Taconic klippe: in De Jong, K.A. , and Scholten, Robert, Gravity
and tectonics: Wiley- Interscience Pubs., New York, New York, pp. 327-342.

and Boucot, A.J., 1965, Origin and distribution of rocks of Ludlow age

(Late Silurian) in the northern Appalachians: Amer. Jour. Sci. , v. 263,
pp. 153-169.

Neuman, R.B., 1967, Bedrock geology of the Shin Pond and Stacyville quadrangles,
Penobscot County, Maine: U.S. Geol. Survey Prof. Paper 524-1, 37p.

, and Rankin, 1966, Bedrock Geology of the Shin Pond Region: iri Caldwell,

D.W., ed. , The Mount Katahdin Region, Maine; New England Intercollegiate
Conf. Guidebook, pp. 8-17.

Noble, J. P. A., 1976, Silurian stratigraphy and paleogeography, Pointe Verte area,
New Brunswick, Canada: Can. Jour. Earth Sci., v. 13, pp. 537-546.

Osberg, P.R., 1968, Stratigraphy, structural geology, and metamorphism of the
Waterville-Vassalboro area, Maine: Maine Geol. Survey Bull. 20, 64p.

20

Pankiwskyj, K..A. ; Ludman, Allan; Griffin, J.R.; and Berry, W.B.N. , 1976,
Stratigraphic relationships on the southeast limb of the Merrimack
Synclinorium in central and west-central Maine: in Lyons, P.C. and
Brownlow, A.H. , eds., Studies in New England Geology: Geol. Soc. Amer.
Mem. 146, pp. 263-278.

Pavlides, Louis, 1965, Meduxnekeag Group and Spragueville Formation of Aroostook
County, northeast Maine: U.S. Geol. Survey Bull. 1244-A, pp. A52-A57.

1968, Stratigraphic and facies relationships of the Carys Mills Formation,

northeast Maine and adjoining New Brunswick: U.S. Geol. Survey Bull. 1264,
44p.

1971, Geologic map of the Houlton quadrangle, Aroostook County, Maine:

U.S. Geol. Survey Geol. Quad. Map, GQ-920.

1972, Geologic map of the Smyrna Mills quadrangle, Aroostook County,

Maine: U.S. Geol. Survey Geol. Quad. Map, GQ-1024.

19 73, Geologic map of the Howe Brook quadrangle, Aroostook County, Maine:

U.S. Geol. Survey Geol. Quad. -Map GQ-1094.

1974, General bedrock geology of northeastern Maine: in Osberg, P.H.,

Guidebook for field trips in east-central and north-central Maine: 1974
New England Intercollegiate Geol. Conf. Guidebook, pp. 61-85.

19 78, Bedrock Geologic Map of the Mars Hill Quadrangle and vicinity,

Aroostook County, Maine: U.S. Geol. Survey Quad. Map 1-1064.

, Mencher, E.; Naylor, R.S.; and Boucot, A. J. , 1964, Tectonic features of

eastern Aroostook County, Maine: U.S. Geol. Survey Prof. Paper 501-C,
pp. C28-C38.

, Boucot, A.J. and Skidmore, W.B., 1968, Stratigraphic evidence for the

Taconic orogeny in the northern Appalachians, in Zen, E-an; White, W.S.;
Hadley, J.B.; and Thompson, J.B., ed. , Studies of Appalachian geology:
Northern and Maritime: New York, Interscience Pubs., Inc., pp. 61-82.

Rast, N. and Stringer, P., 1974, Recent advances and the interpretation of

geological structure of New Brunswick: Geoscience Canada, v. 1, pp. 15-25.

Richter, D.A. and Roy, D.C., 1974, Subgreenschist metamorphic assemblages in
northern Maine: Can. Min. , v. 12, pp. 469-474.

, and Roy, D.C., 1976, Prehnite-pumpellyite facies metamorphism in central

Aroostook County, Maine: in Lyons , P.C. and Brownlow, A.H. , eds., Studies
in New England geology, Geol. Soc. America Mem. 146, pp. 239-261.

Rodgers, John, 1970, The tectonics of the Appalachians: New York, Wiley and
Sons, 271p.

Roy, D.C., 1970, The Silurian of northeastern Aroostook County, Maine (Ph.D.
thesis): Cambridge, Massachusetts Institute of Technology, 483p.

21

19 73, The provenance and tectonic setting of the Frenchville Formation,

northeastern Maine: Geol. Soc. America Abs. with Programs, v. 5, p. 214.

, and Mencher, E. , 1976, Ordovician and Silurian stratigraphy of north-
eastern Aroostook County, Maine, in Page, L.R., ed. , Contributions to
the stratigraphy of New England: Geol. Soc. America Mem. 148, pp. 25-52.

19 78, Geologic Map of a Portion of Northeastern Aroostook County, Maine:

Maine Geol. Survey Open File Map.

Ruitenberg, A. A. and Ludman, Allan, 1978, Stratigraphy and tectonic setting of
early Paleozoic sedimentary rocks of the Wirral-Big Lake area, south-
western New Brunswick and southeastern Maine: Can. Jour. Earth Sci., v. 15,
pp. 22-32.

Schopf, J.M., 1964, Middle Devonian plant fossils from northern Maine: U.S.
Geol. Survey Prof. Paper 501-D, pp. D43-D49.

Skinner, Ralph, 1974, Geology of Tetagouche Lakes, Bathurst, and Nepisiquit
Falls Map-areas, New Brunswick: Geol. Survey of Can. Mem. 371, 133p.

St. Julien, Pierre and Hubert, Claude, 1975, Evolution of the Taconian orogen
in the Quebec Appalachians: Amer. Jour. Sci., V.275-A, pp. 337-362.

Thompson, J.B. and Norton, S.A. , 1968, Paleozoic regional metamorphism in New
England and adjacent area, in Zen, E-an; White, W.S.; Hadley, J.B.; and
Thompson, J.B., eds., Studies of Appalachian geology: Northern and
Maritime: New York, Interscience Pubs., Inc., pp. 319-327.

Venugopal, D.V. , 1979, Geology of Debec Junction-Gibson, Millstream-Temerance-
ville, Meductic Region; Map-areas G-21, H-21, 1-21, and H-22; New
Brunswick Department of Natural Resources, Minerals Resources Branch,
Map Report 79-5 Venugopal, 19 79.

White, K.J., 1975, The Mapleton Formation: An immediately post-Acadian basin
fill (MS thesis): Chestnut Hill, Ma., Boston College, 70p.

22

OUTLINE OF THE PLEISTOCENE GEOLOGY OF NORTHERN MAINE

AND ADJACENT CANADA
by
Andrew N. Genes
Department of Regional Studies
Boston State College
Boston, Massachusetts 02115

INTRODUCTION

For some time, questions concerning the intensity, style, and limits
of glaciation throughout northern New England, adjacent Canada, and the
Maritime Provinces have been the subject of much controversy.

Initially, Chalmers (1895) put forth the concept of glaciation based
on the influence of local centers of ice accumulation. These local ice
centers were to have coalesced at maximum glaciation but were thought to
have still remained partially independent. As investigations continued
this view remained viable and is still considered tenable (Prest and
Grant, 1969; Grant, 1977) although modified with respect to specifics
concerning limited isostatic rebound, areas showing limited or no traces
of glaciation, and stratigraphy.

Alternatively, because of characteristics and features related to
strong regional ice flow, lithologic transport, and alpine glaciation,
Goldthwaite (1924) advanced the concept of continental glaciation which
was controlled by southward flowing lobes of ice emanating from a massive
continental ice sheet.

This latter view has been expanded to a more sophisticated level by
using modern day analogues such as Antarctic (Hughes, 1973; Boms, 1979,
pers comm.) and Greenland (TenBring, 1974) glaciers, along with concepts
of calving bays (Thomas, 1977), ice streams (Hughes et al., 1977) and,
more recently, models employing the concept of the thermal regime of
ice masses (Sugden, 1977; Hughes, 1979, pers. comm.).

23

Currently, both perspectives are receiving renewed stimulus.

With the exception of the classic work of Leavitt and Perkins (1935)
northern Maine has only recently begun to be mapped on a reconnaissance
basis, primarily because in the past, access was difficult in what were
generally regarded as "wilderness" areas. However, land use planning
considerations in conjunction with accelerated logging operations led
to various programs starting several years ago involving reconnaissance
surficial mapping (Prescott, 1973a; 1973b). It was primarily through
these programs, fostered and sponsored by the Main Geological Survey,
that northern Maine became an active area of Pleistocene geological
investigation. Fortunately so, for prior to these programs most Pleisto-
cene investigations were concentrated along southern and coastal Maine
leaving a large, unmapped void between that area and Canada, x^here
Canadian colleagues were working.

As it turns out, this region of northern Maine and adjacent Canada
seems destined to play a pivotal role in the deciphering of Pleistocene
events along the northeastern corridor of the United States and Canada.

Areally, northern Maine consists of the largest counties in the
state; Aroostook County (the largest) forms the northwest, north, and
eastern borders of Maine as it abuts Quebec and New Brunswick, separated
from Canada to the north by the St. John River (international boundary);
Franklin and Somerset Counties delimit the western border of Maine where
they join the Quebec border which is roughly coincident with the Boundary
Mountains; Piscataquis and Penobscott Counties consititute the central -
southern region of northern Maine.

For the most part, the region is underlain by the cyclically bedded
grey slate and metasandstone Seboomook Formation. Some large areas of
metamorphosed volcanic rocks occur scattered throughout the region. In
central Piscataquis County biotite - muscovite granites and quartz
monzonites outcrop, increasing in occurrence to the south. Metamorphosed
sandstone, silts tones, and limestones outcrop near the eastern Maine-New
Brunswick border (See Figures 1 and 2 of Roy, Trip B-6, this vol-
ume)

24

Topographically, the area is a dissected upland plateau with
regional bedrock structures trending northeasterly to southwesterly,
transverse to knox<m ice-flow direction.

STRATIGRAPHY AND EVENTS

The stratigraphy of glacial deposits of southern Quebec and the
events which they represent (Table 1) is based on work by McDonald and
Shilts (1971). They determined that the last glacial advance in southern
Quebec is represented by the Lennoxville Till, having a northwest pro-
venance, and that southeastern Quebec was deglaciated by about 12,500
years BP. .

This Lennoxville Till is correlated with the surface till of Maine
(Borns and Calkin, 1977). Coastal Maine was covered by ice 13,200 years
BP.(Schlee and Pratt, 1970) and readvances to the coast occurred during
general ice recession (Borns, 1966; 1973). An active ice margin existed
along the eastern coast of Maine (Borns, 1966) approximately 13,300 years
BP.. At about the same time an active ice margin existed in southwestern
New Brunswick (Gadd, 1973). Both McDonald and Shilts, and Borns and
Calkin consider the Lennoxville Till as representing the entire late
Wisconsinan time interval. Borns and Calkin (1977) suggest that the
final deglaciation of Maine occurred by thinning and stagnation of ice
throughout the uplands of northwestern Maine with no reorganization of
flow to form an active center of ice dispersal. In support of this
contention, they suggest that cirque glaciers did not develop in the
mountains of west-central Maine following dissipation of the ice. This
implies that the regional snow line had become elevated above the highest
mountains prior to their emergence. This view is supported by Davis
(1976) but is in opposition to views held by Thompson (1960, 1961a, 1961b)
and Caldwell (1966) .

Borns and Calkin (1977) visualize that by the time the ice margin
had retreated to the proximal side of the coastal moraine belt all

25

Time-Stratiaraphic
Unit"

Rock-Stratigraphic Unit

Chronologic Control

C71

a

IS)

c

in

e
o
o

3

01

+J

post-Lennoxville Sediments

12,640+190 - peat (GSC-312)
12,570+220 - peat (GSC-419)
12,000±230 - marine shells (GSC-936)
11,500+160 - marine shells (GSC-475-2)

ai

•o
•a

i

Lennoxville Till

Gayhurst Formation

>20,000 B.P. (GSC 1137)
Ca 4000 varves

Chaudiere Till

>>

■a

01

Massawippi Formation

>54,000 B.P. (Y-16G3)
> 41, 000 B.P. (GSC-507)
>40.000 B.P. (GSC-1084)

Johnville Till

pre-Johnville Sediments

Table I. Quaternary Stratigraphic Column, Southeastern Quebec; (after McDonald and Shilts,
1971)

26

Canadian ice flow was southward within the active ice mass north of the
Boundary Mountains. During this period of stagnant ice in Maine, the
Canadian ice constructed successive frontal moraines as the ice margin
receded to the St. Lawrence lowland of Quebec (Gadd, 1964; McDonald and
Shilts, 1971).

Glacial Lake Madawaska was formed in New Brunswick when the St.
John River at Grand Falls, New Brunswick was dammed by the Grand Falls
Moraine (Lee, 1955). Recession of the ice front from Grand Falls had
to have occurred well before 10,200 years BP. because peat overlying till
at Green River, New Brunswick has been dated at 10,200 + 350 years BP.
(Lee, 1955). Kite (1979) has had this peat redated and concludes that
it possibly could be as old as approximately 12,000 years BP. Until
recently (Genes and Newman, 1979* Brewer, 1980) end moraines had not been
reported between the end moraine belt around Pineo Ridge, at coastal
Maine, and the international border at the west side of the Boundary
Mountains. Therefore, the limitations imposed by previous investigations
of adjacent areas required the conclusion that all of the surface till of
Maine be Lennoxville in age, that is, it is a simple remnant of the last
major advance to the coast. In turn, this implied that subsequent to
the invasion of the Champlain Sea, approximately 13,000 years BP.,
Laurentide ice did not transgress into Maine (Lasalle et al., 1977) or
even at all during late Wisconsinan time (Grant, 1977).

It must be mentioned that there are extremely divergent views as to
the behaviour of the ice in southeastern Canada. Gadd (1973) envisions
the Highland Front Moraine system in Quebec as having been emplaced as a
recessional moraine during normal retreat of the margin of southerly
flowing ice. Lasalle et al. (1977) visualize the system as having
formed from northerly flowing ice which resulted from the division of the
Laurentide ice sheet into separate lobes by the Champlain Sea embayment.
Grant (1977) because of his adherence to the concept of Appalachian ice
centers maintains that Laurentide ice never reached Maine.

Field mapping in northern Maine and southern Canada now implies
that deglaciation was more complex than previous work has suggested.

27

Evidence from striations, moraines, outwash overrun and capped by till,
and three or possibly four distinct tills can be interpreted as having
resulted from multiple glaciations: from different flow regimes within
a single ice sheet; or from penecontemporaneous deposition from different
ice centers .

The following considerations are implicit in any interpretation
regarding the late Wisconsinan history of northern Maine and adjacent
Canada:

1) A late Wisconsinan readvance formed the Pineo Ridge
Moraine along the eastern coast of Maine at approximately
12,700 years BP. (Boms, 1966). The minimum age of de-
glaciation of southwestern New Brunswick is implied to

be 12,600 ± 279 years BP. (Gadd, 1973).

2) The date of 12,600 years BP. for the Highland Front
Moraine (Gadd, 1964) in Quebec, represents the minimum
date for the incursion of the Champlain Sea into the
St. Lawrence Valley. This date also represents the
minimum date at which the Laurentide ice became de-
tached from the Laurentide ice north of the St. Law-
rence River.

3) All observed evidence indicates that northern Maine was
overrun by ice moving in a southeasterly direction, and
that glacial recession was accomplished by downwasting
and recession from coastal to northern Maine. Evidence
in northwestern Maine indicates ice flow to the north.

In light of these controversies, this years' surficial field trips should
prove interesting indeed.

Claude Gauthier has been deciphering the New Brunswick Pleistocene
for several years and will lead a trip on the New Brunswick side of the
St. John River Valley. Glacial stratigraphy around Edmunston, a

28

reappraisal of Glacial Lake Madawaska, and sections of the Grand Falls
Moraine complex, will be discussed. (Trip B-9)

Kite and Borns will discuss deposits associated with Glacial Lake
Madawaska as they relate to both sides of the middle reaches of the St.
John River. Although, primarily concerned with Holocene events, they
suggest important concepts regarding the determination of deposits and
how they should be mapped. (Trip C-4)

Wisconsinan stratigraphy along the southern bank of the St. John
River, including multiple till sections which permit interpretation of
late Wisconsinan deglaciation in northern Maine, is the subject of the
trip by Genes and Newman. (Trip B-8)

Genes, Newman, and Brewer will examine moraines, eskers , and other
landforms in northern Maine which suggest the mode and extent of till
emplacement in that region. (Trip C-6)

D. Caldwell will review the Wisconsinan alpine glaciation of Mt .
Katahden. In case a climb is not possible because of inclement weather,
a surficial trip of the area is planned. (Trips A-3, B-2)

29

REFERENCES CITED

Borns, H.W.Jr., 1966, An end Sforaine complex in southeastern Maine

(Abstract): Geol. Soc. America Abs . , Special Paper 101, p. 13-14.

, 1973, Late Wisconsin fluctuations of the Laurent ide ice

sheet in southern and eastern New England, _in The Wisconsinan
Stage, Black, R.F. and others, editors: Geol. Soc. America Mem. 136

, and Calkin, P.E., 1977, Quaternary glaciation, west cen-
tral Maine: Geol. Soc. America Bull., v. 88, p. 1773--1784.

Brewer, T.B. , 1980, Surficial geology of the Amity Quadrangle, Aroostook
County, Maine: Open File Report, Maine Geol. Surv. , Augusta, Maine

Caldwell, D.W. , 1966, Pleistocene geology of Mt. Katahdin, in Caldwell,
D.W., ed., N.E.I.G.C. Guidebook: p. 51-61.

Chalmers, R. , 1895, Surface geology of eastern New Brunswick, north
western Nova Scotia and a portion of Prince Edward Island:
G.S.C. Annual Rep., v. 7, p. 143.

Davis, P.T., 1976, Quaternary glacial history of Mt . Katahdin, Maine:
(M.S. Thesis): Orono, Univ. Maine, p. 155.

Gadd, N.R., 1964, Moraines in the Appalachian region of Quebec: Geol.
Soc. America Bull., v. 75, p. 1249-1254.

-, 1973, Quaternary geology of southwest New Brunswick with par-

ticular reference to Frederic ton area: Geol. Surv. Can., Dept. of
Energy, Mines and Res., Paper 71-34.

Genes, A.N., and Newman, W.B., 1979, Late Wisconsinan glaciation of
Aroostook County, Maine: Geol. Soc. America Abst., Northeast
Section, p. 36.

30

Goldthwaite, J.W., 1924, Physiography of Nova Scotia: G.S.C. Memoir 140,
p. 212.

Grant, D.R., 1977, Glacial style and ice limits, the Quaternary strati-
graphic record, and changes of land and ocean level in the Atlantic
Provinces-Canada: Geog. Phys . and Quat . , v. 31, no. 304, p. 247-260

Hughes, T., 1973, Glacial history of the Antarctica ice sheet: Jrn.
Geophys. Res., v. 78, p. 7884-7998.

, Denton, G.H., and Grosswald, M.G. , 1977, Was there a late Wurm

Arctic ice sheet?: Nature, v. 266, p. 596-602.

Kite, J.S., 1979, Postglacial geologic history of the middle St. John

River Valley: Unpublished M.S. Thesis, University of Maine, Orono.

Lasalle, P. Martineau, G. , and Chauvin, L. , 1977, Morphology stratigraphy,
and deglaciation in Beauce-Notre Dame Mountains-Laurentide Park
area: Min. of Nat. Res., DPV-516.

Leavitt, H.W. , and Perkins, E.H., 1935, Glacial geology of Maine: v. 2,
Maine Tech. Expt. Sta. Bull. 30, p. 232.

Lee. H., 1955, Surficial geology of Edmunston, Madawaska, and Temiscouata
Counties, New Brunswick and Quebec: Geol. Surv. of Canada,
Paper 55-15.

McDonald, B.C. and Shilts, M.W., 1971, Quaternary stratigraphy and events
in southeastern Quebec: Geol. Soc. America Bull., v. 82,
p. 683-698.

Prescott, G.C.Jr., 1973A, Groundwater favorability and surficial geology
of parts of the Meduxnekeag River and Prestile Stream basins: Hyd.
Inv. Atlas HA-486, U.S.G.S..

31

, 1973B, Groundwater favorability and surficial geology

of the lower St. John River Valley, Maine: Hyd. Inv. Atlas HA-485,

Prest,V.K., and Grant, D.R., 1969, Retreat of the last ice sheet from the
Maritime Provinces, Gulf of St. Lawrence Region: G.S.C. Paper 69-
33, p. 15.

Schlee, J., and Pratt, R.M., 1970, Atlantic continental shelf and slope

of the United States, gravels of the northeastern part: (U.S.G.S.)
Prof. Paper 529 H, p. 39.

Sugden, D.E. 1977, Reconstruction of the morphology, dynamics, and ther-
mal characteristics of the Laurentide ice sheet at its maximum:
Arctic and Alpine Res., v. 9, p. 21-47.

TenBring, L.W., 1974, Glacio-isostasy and glacial history, western
Greenland: Geol. Soc. America Bull., v. 85, p. 921-943.

Thomas, R.H., 1977, Calving bay dynamics and ice sheet retreat up the

St. Lawrence Valley system: Geog. Phys . and Quat., v. 31, p. 347-356

Thompson, W.F., 1960, The shape of New England mountains, Part I:
Appalachia, v. 33, p. 145-159.

-, 1961a, The shape of New England mountains, Part II:

Appalachia, v. 33, p. 316-335.

, 1961b, The shape of New England mountains, Part III

Appalachia, v. 33, p. 458-478.

32

TRIP A-l

GEOLOGY OF THE BOTTLE LAKE COMPLEX, MAINE

Robert A. Ayuso and David R. Wones
U.S. Geological Survey and
Virginia Polytechnic Institute and State University

Introduction

The Bottle Lake Complex consists of two late Paleozoic, massive
granitic plutons exposed in an area of about 1100 km in the Scraggly Lake,
Springfield, Winn, Wabassus Lake, Nicatous Lake, and Saponac 15-minute
quadrangles, and located between U.S. Rts. 6 and 9 in east-central Maine
(Figure 1). The Complex comprises two overlapping, subcircular, generally
east-west trending bodies intruding greenschist facies metamorphic rocks
of the Merrimack Synclinorium. The bodies have a granitic extension striking
to the northeast (Figures 1, 2).

Larrabee and others (1965) mapped the granitic and metamorphic rocks
in the Big Lake region in reconaissance. Their study of the granitoid
rocks represented by the Bottle Lake Complex made no field or petrographic
distinctions between these rocks except for the possibility of an internal
contact between the main body of the Complex and the large northeast-trending
granitic protuberance (Topsfield facies). Observed field and petrographic
differences were ascribed to secondary processes, mainly as a result of
country-rock assimilation. However, Ayuso (1979) recognized several map-
pable petrographic types that exhibited a consistent arrangement and in-
dicated the presence of two separate intrusive granitoid plutons. From
east to west, the two granitoid plutons consist of the granite of Whitney
Cove and the granite of Passadumkeag River, respectively. The Topsfield
facies is assigned to the Whitney Cove pluton because of similarity in
their petrographic and field relations and absence of internal contacts
(Figure 2).

Metamorphic rocks in the region are presently under intensive
scrutiny by A. Ludman (1978a, 1978b) and coworkers in the area to the
northeast and east of the Complex, and by Wones (1979) in the Norumbega
fault region to the south. Detailed mapping by Olson (1972) to the south-
west of the Complex and the reconaissance work of Doyle and others (1961)
and Cole (1961) concentrated respectively, to the southwest and north.
As summarized by Ludman (1978b) , metamorphic rocks in this area are in
the greenschist facies and exhibit variable age and lithology (Figure 2).
These rocks range in age from Cambro-Ordovician (?) to Siluro-Devonian,
and from almost monomineralic sandstones to andesitic volcanics. Most
traverses from country-rock toward the granite show well-developed min-
eral zonation as a result of the contact metamorphic effects (Ayuso, in
press) .

33

45°

68°

\

fi?

1 ^

*<&$*

0*w

f&

44°

Figure 1 - State of Maine showing the location of the Bottle
Lake Complex. Light stipple - granite and granodiorite; dark
stipple - diorite and gabbro. 1-Lucerne pluton; 2-Lead Moun-
tain pluton; 3-Bottle Lake Complex; 4-Center Pond pluton; 5-
Seboeis foliated granitoid; 6-Seboeis non-foliated granites;
7-Katahdin pluton. NFZ-Norumbega Fault Zone. (from Loiselle
and Ayuso, 1980).

34

•H
to

a

P<

8

O

CO

o

PQ
<D

o

a

o

•H

o
o

0)
00

•o

0)
N
•H
iH
03
M
0)

§

CN

CU

U

00
•H

35

PLUTONIC ROCKS - Bottle Lake Complex

Dpc

Granite of Passadumkeag River

Coarse-grained biotite-hornblende bearing facies - generally
quartz monzonitic composition, high color index, euhedral
amphibole prisms, and abundant mafic inclusions

Coarse-grained biotite-hornblende granitic to quartz-monzonitic
facies - commonly heterogeneous in texture; quartz forms large
mosaic textures, generally poorer in mafic minerals than rocks
in the interior

Dprb

Amphibolite

Dd

Diabase (?) intruded in cataclastic zone

Granite of Whitney Cove

Dgwp

Dgw

Medium to coarse-grained, biotite bearing facies - homogeneous in
composition, typically porphyritic, and low color index

Coarse-grained biotite facies granite - equidimensional texture,
low color index, biotite in pseudohexagonal books

PLUTONIC ROCKS - Other Granitic Rocks

Dgc

Dglm

Dwl

Granite of Center Pond - medium-grained, granitic rock ranging
from diorite to biotite granite; contains abundant hornblende

Undifferentiated granitic rocks, usually coarse-grained, biotite-
hornblende bearing rocks

Medium-grained granite found only within Norumbega fault zone,
lower mafic-mineral content than Bottle Lake Complex

METAMORPHIC ROCKS

DSv

SOv

SOp

COu

Vassalboro Fm (?) - calcareous silt stones and pelites showing
prominent mineralogic zonation from contact metamorphic effects
imposed by the Bottle Lake Complex

Undifferentiated Siluro-Ordovician (?), volcanic rocks, rusty
pelitic beds

Undifferentiated Siluro-Ordovician (?), pelitic siltstones
containing graded beds

Graywakes, siltstones, slates are the main lithologies

Geology by R. A. Ayuso and assistants (1977,1978,1979); based on fieldwork by
D. Larrabee and assistants (1965), A. Ludman (1978a, b), D. R. Wones (1979),
and Scambos (1980).

36

Field Relations

Field and petrographic contrasts between the Whitney Cove and
Passadumkeag River plutons include their strikingly different textures
and mineralogy, particularly the kind and abundance of ferromagnesian
phases, abundance of aplites and pegmatites, and abundance and type of
inclusions. Field relations are summarized in Ayuso (in press). With
the exception of amphibole, both granites are generally coarse-grained,
consist of two feldspars, biotite and quartz, with minor amounts of
primary sphene, zircon, allanite, apatite, oxides (magnetite and ilmenite)
and sulfides. Sharp internal contacts, dikes and inclusions of older in
younger granite are notably absent, but the relative sequence of intrusion
may be established from mapping petrographic types, cross-cutting relation-
ships of faults, and preliminary age determinations.

The granite of Whitney Cove

2
The granite of Whitney Cove covers an area of about 400 km . Its

mineralogy is consistently granitic (Streickeisen, 1973), has relatively

low total ferromagnesian mineral content, lacks amphibole and shows either

a predominant seriate (rim facies) or porphyritic (core facies) texture.

Aplites and pegmatites are generally common in most exposures, while mafic

inclusions are smaller and less common than in the Passadumkeag River

pluton.

This pluton consists of three units: 1) Topsfield facies, 2) rim
facies, and 3) core facies. The generalized geologic map shown on Figure
2 does not separate the Topsfield facies from the rim facies of the Whitney
Cove pluton because of their similarity in field relations and absence of
an internal contact between them. The Topsfield facies is commonly a
medium to coarse-grained, reddish rock of low color index, approximately
400 m.y. old as determined by U-Pb work on zircons (R. Zartman, pers.
coram. ) . Away from the main body of the pluton of Whitney Cove, it be-
comes finer-grained, progressively pinker, lower in color index and more
intensely sheared. The rim facies is a pink, medium to coarse-grained,
hypidiomorphic, seriate rock of low color index. No mineralogical
characteristics distinguish the core from the rim, but textural dif-
ferences are, however, significant and may be used to delineate the ex-
tent of each; contacts between the two are for the most part gradational.
Within the core rocks, biotite is sometimes present as phenocrystic clots,
but it is more abundant as fine-grained aggregates disseminated through
the rock. The predominant phenocrysts are alkali-feldspars (up to 3.5 cm)
and these are usually euhedral but also show edges embayed by biotite,
plagioclase and quartz; some of the alkali-feldspars show rapakivi textures.
Plagioclase mantled by alkali-feldspar (anti-rapakivi texture) is rarely
present in these rocks. Quartz is usually constrained to the matrix, how-
ever, it is also present as a phenocryst (up to 1 cm) phase. Biotite is
characteristically enclosed within subhedral, phenocrystic (up to 2.5 cm)
plagioclase; less commonly, biotite forms euhedral phenocrysts similar to
those found in the rim facies. Fine-grained biotite and plagioclase make
numerous clusters in the matrix and give this rock an easily recognizable
speckled appearance.

37

The consistently granitic composition of the Whitney Cove pluton is
shown in Figure 3, where the modal variation obtained by counting stained
slabs is displayed. The quartz-plagioclase-alkali feldspar ternary shows
complete overlap between core and rim facies, and this is reinforced in
the accompanying ternaries between mafic and felsic minerals. Also note
the low scatter in total ferromagnesian content and the absence of a clear
concentration gradient of mafic-poor to mafic-rich rocks from rim to core
(Figures 4,5). Both facies have approximately 4% total mafic mineral
content. The general mineralogical homogeneity of the pluton is further
shown by its lack of correlation between alkali-feldspar and mafic-mineral
content in Figure 5.

The granite of Passadumkeag River

The two facies in the pluton of Passadumkeag River are the rim and
core (Figure 2). Generally pink rocks of seriate to equidimensional
texture, finer-grained and lower in color index characterize the rim.
Amphibole is not abundant but it is usually present, at least in trace
quantities. The low content of ferromagnesian phases and many of the
textures are reminiscent of the rim facies of the Whitney Cove granite.
Preliminary Rb-Sr isotopic determinations on whole-rock samples suggests
an age of about 360 + 16 m.y.

In contrast, rocks found in the interior of the pluton are invariably
mafic-rich, commonly showing abundant, euhedral amphibole, and numerous
mafic-schist, metasedimentary, and quartz-dioritic inclusions. Charac-
teristically, this facies consists of gray porphyritic rocks with prominent
subhedral to euhedral alkali feldspar phenocrysts (3 to 5 cm) that are
usually embayed and mantled by plagioclase in a typical rapakivi texture.
Many of these phenocrysts contain inclusions of euhedral plagioclase, quartz,
and ferromagnesian minerals which help to delineate several growth zones.
Euhedral phenocrysts of plagioclase (3 to 5 cm) are commonly filled with
inclusions; together with fine-grained ferromagnesian minerals, plagioclase
forms clots that alternate with quartz-rich clusters in the matrix. A
common characteristic of the Passadumkeag River granitoid is that euhedral
amphibole (0.5 cm) commonly encloses biotite.

In contrast to the Whitney Cove pluton, marked variations exist
between the rim and core facies of the granite of Passadumkeag River
(Figure 3). In this case, a granitic rim envelops a quartz monzonitic
core of consistently higher plagioclase and mafic-mineral content. As
shown in Figure 4, the core facies has almost double the mafic-mineral
content of the rim; this zonation is clearly exhibited in the reversed
correlation of alkali-feldspar and mafic content between rim and core
(Figure 5).

The Whitney Cove pluton generally contains more felsic dikes
(aplites, pegmatites, granophyres) than the granite of Passadumkeag
River. These dikes are typically less than 0.7 cm wide, exhibit diffuse
contacts, and are concentrated near the granite-country rock and granite-
granite contact.

38

MAFICS

THE BOTTLE LAKE COMPLEX

QUARTZ

MAFICS

PLAGIOCLASE

FELDSPAR

GRANITOID OF PASSADUMKEAG RIVER

* CORE
a RIM
+ NORTHEAST EXTENSION

GRANITOID OF WHITNEY COVE

• CORE

° RIM

MAFICS

Figure 3 - Modal mineralogy of the Bottle Lake Complex. North-
east extension refers to the granitic rocks of the Passadumkeag
River pluton intruded north of the cataclastic zone.

39

Core 0 x = 4.7

Rim D x = 4.2

x = 4.3

Granitoid of Whitney Cove

20

16
12
8

oL

s^n __

I 5 9 13 17 21
Biotite + Hornblende

f 10

Granitoid of Passadumkeag
River

Core ^ x = 10.6
Rim □ x = 5.7
x = 8.6

m

5 9 13 17 21
Biotite + Hornblende

Passadumkeag
River

H Whitney Cove

5 9 13 17 21

Biotite + Hornblende

Figure 4 - Histogram of the total abundance of mafic minerals in the
Bottle Lake Complex

40

X

UJ

o
o

UJ

<

UJ

-J
o

GO
UJ

CD

>
O

o

>%

CD

c

o

"O

o

c
o

o

CD

or o

+ +

+"1

+

8.*r

^ o

i

i

lo

CD

■o

jCD

o

X

o
CD

o

lO

o

o
ro

O
t\J

O

jDdsp|3-j - y\

c
o

<*

k_

CO

CD

c

—

>

CD

cr

X

LU

o>

o

-*—

CD

CO

O

«

E

M/

< <

3
■o
o

CO

£ ~ *-

— O O

tr o z

o
Q_

14—

« « +

Granitoid o

i-

+
<

+

<

i i i

o

c\J

CD

2 +

CD

- LO

o
GO

O

CD

O

LO

o

o

rO

o

c\J

JDdsp|8-J - >j

x
a)

e -h

o

o c
o

0) 4-)

^! 3

Cd tH

t-J ex

a) }-i

4J >

4-> «H

O Ctf

too

cu cd

,c a)

4-> ^

<4H 3

O T3

cd

4J co

C co

a) cd

4-> PL,

C

O CU

Cd M-(

^ o
cu

C CO
•H 4*$

6 ^
o

O S-i
•H

m-i a

•H
4-1
•H

c

CO Cd •
> U CU

}-i O

cd <U N

CO 4-1

cd

6

■*- cd

o

•H
4J
CO

cd

CO .— I

5-i O

iH a) cd

Cd U-* 4-1

^ cu cd
5-1 o

a) 4-t

i

•H

C CU

m-i o x:

O 'H 4->
CO

4J C M-l

o cu o

I

pi* H Xi

CU 4-1
S-i
4-1 O

m co c
cd

CU CU T3
5-i X CU

3 w-d

00 5-i 3
•H O 5-i
U* Z 4->

41

An amphibolite unit of limited length outcrops between the two
plutons (Figure 2), and it is concentrated near the fault zone that cuts
the Whitney Cove pluton. Typical rocks associated with this unit range
from black to green in color and contain prominent orange feldspar pheno-
crysts; these rocks are intimately mixed with granitic rocks of the Whitney
Cove pluton. Aplites, pegmatitic pods and large inclusions of variable
lithology (from felsic to mafic mineralogy) are commonly found within the
amphibolite unit.

Structures

The Bottle Lake Complex is generally massive in fabric, often even at
the granite-country rock contact. Several areas within the Complex show
well-developed foliation. These areas may be grouped as follows:
1) randomly distributed zones in the core facies of the Passadumkeag River
pluton; 2) the area associated with the amphibolite rocks near a zone of
cataclastic deformation; and 3) the randomly foliated rocks in the core
of the Whitney Cove pluton. The c_ dimensions in the feldspars are aligned
parallel to the northeast in the core rocks of the Passadumkeag River
granite and in the amphibolite rocks. In contrast to this, no discernible
preferred orientation in the regional sense is apparent in the foliated
rocks found in the core of the Whitney Cove pluton.

Although joints and fractures are common in the Complex, they typically
show great scatter even within individual outcrops. The only exception to
this lack of preferred alignment is found in those joint sets associated
with cataclastic zones.

Three cataclastic zones are present in the Bottle Lake Complex:
1) the Norumbega fault zone, 2) an unnamed EW fault, south of the Topsfield
facies and exposed near U.S. Rt. 6, and 3) the NE-trending belt that cuts
the Whitney Cove pluton (Figure 3) . The Norumbega fault system forms the
southern contact of the Whitney Cove granite. According to Wones (1979),
the amount of displacement along the fault is unknown. The EW fault is
exposed near U.S. Rt. 6 between the main mass of the Whitney Cove granite
and the Topsfield facies and shows left-lateral displacement (Ludman, 1978b).
Exposure of this fault is expressed by sheared, epidotized, red rocks typ-
ical of the Topsfield facies. Finally, the NE-trending cataclastic zone
through the pluton of Whitney Cove forms a wide band (1.5 km?) which termi-
nates against the Passadumkeag River granite.

Bulk Chemistry

Preliminary major and trace element X-ray fluorescence analyses of the
Bottle Lake Complex are shown in Figure 6. In spite of the significant
mineralogical contrasts between the two plutons, the data suggests a re-

42

l \)

■

875,

?

275

-

^

-

|225

-

Q.

_

****

in ,75

-

125

-

■

75

25

>

O II

/

CJ

+ . 9

CM >

■

o ^ -,

2 7

/

13

L - '

r//r

-i — i — i — n

■%-

O

o

O

O

1.0
0.6
0.2

^ 22

0s*

•t 18

O

_w 14 -
<

Passadumkeag River

* Core
A Rim

° Mafic inclusions

Whitney Cove

• Core
o Rim

Metamorphic rocks

-//-

-//-

□

53

V^^

1

^

— -ic>#

PA.^^,

J L

59 62 65 68 71

Si02 ( wt. % )

74

77

Figure 6 - Silica variation diagrams of selected oxides and
strontium of the Bottle Lake Complex. Stippled field re-
presents analyses of the Lucerne and Lead Mountain granites;
field outline represents the Center Pond pluton (Scambos, 1980)

43

markable colinearity in the trend of both granitoids, showing similar
ranges in SiCL content, and complete ovelap between data for samples
from the two cores and rims. As expected, the core of the Passadumkeag
River granite is less silica-rich and more mafic in character than its
rim; the pluton shows reverse zonation. However, the Whitney Cove core
facies, despite its mineralogical similarity to the rim, is extremely un-
like it in composition, and instead resembles the core of the Passadumkeag
River pluton. Reverse zonation is chemically expressed in both plutons,
but it is mineralogically evident only in the Passadumkeag River granitoid.

Strontium concentrations reinforce the contrast between core and rim
of these plutons (Figure 6). As in the major element trends, a reverse
zonation is clear as cores are substantially enriched in strontium com-
pared to the rims. Unlike the major element plots, however, a difference
is suggested in the core of the Whitney Cove pluton which consistently
shows higher strontium content compared to the Passadumkeag River core.
An additional suggestion from this data is, in spite of the overlap, that
each pluton describes parallel but distinct evolution trends from strontium-
rich to strontium-poor rocks.

Mineral Chemistry

Biotite is the most common f erromagnesian phase in the Bottle Lake
Complex; in spite of the different petrographic character among the rocks
of the Complex, no consistent differences are expressed in its optical
properties. However, on the basis of microprobe analyses of represent-
ative samples from each granitoid, biotite shows distinct compositional
variations and trends. In general, biotite in the Whitney Cove pluton
is lower in Fe/(Fe + Mg) ratio (0.55 to 0.65) than that in the Passa-
dumkeag River granite (0.60 to 0.80). Also, biotite from Whitney Cove
tends to have higher F, but lower TiO„^ Figure 7 shows that both plutons
have decreasing Ti with increasing Al and overlap between cores and
rims. Biotite in the Passadumkeag River pluton shows a larger range in
Al compared to that in Whitney Cove; near the granite-country rock con-
tact, it tends toward extreme enrichment in Al

Amphibole is concentrated within the Passadumkeag River pluton and
generally shows a lack of correlation in traverses from the edge of the
core toward the interior. As in biotite, amphibole shows a limited range
in Fe/(Fe + Mg) from 0.55 to 0.75 without a distinct range within each
facies. Figure 8 shows that amphibole compositions tend to increase in
Al with increasing A site content. A prevailing feature of the Passa-
dumkeag River pluton is the common intergrowth and inclusion of biotite
in amphibole. According to the studies of Wones and Dodge (1977) on po-
tassic magmas, crystallization of biotite prior to amphibole is facil-
itated by water deficient conditions.

Preliminary work on feldspar compositions show that plagioclase ranges
from pure albite to An for the Complex as a whole. Most plagioclase
cores are usually richer in anorthite, show normal zonation, and are en-
veloped by strongly zoned albitic rims. Between the two plutons, the Whitney

44

BIOTITE

0.26 -

0.22 -

0.18 -

0.14

0.24

1 1 1 1 1 1 1 1 1 1

▲

A A

Passadumkeaq River

± Core Facies
A Rim Facies
1 1 1 1 1 1 1 1 1 1

1 1 1 1 1

*>*

A

1 1 1 1 1

1 1 1

A

1 1 l

1 1 1
i

i

ill-
i i i

i

i

0.20-

0.16 -

0.00

1 1 1 1 1 1 1 1

1 1 1

1 1 1

1 1 1

1 1

1 1 1

1

o o

• #

— '

o

•

og>

•

—

O Q>

O
O

o

o

•

—

o

o

—

Whitnev Cove

• Core Facies

—

O Rim Facies

i i i i i i i i

1 1 1

1 1 1

1 1 1

1 1

1 1 1

1

0.05

0.10

0.15

0.20

0.25

Al

m

Figure 7 - Representative microprobe analyses of biotite from the
Bottle Lake Complex showing cation proportions of Ti and Al^I in
one-half of the formula unit. The stippled field represents
biotite near the granite-country rock contact. Error in deter-
minations shown approximately at 2o level.

45

-^ Lj

CO

<

I-
o

e

o

4->

c

3

•H

iH

&.

U
O

U

M

CU

H

>

w

•H

Pd

60

4J

cd

•H

cu

c

M

3

CO

T3

«H

cd

0

CO

e

CO

n

CO

o

Ph

<4-l

CU

CU

X

42

4J

4-1

c

<4-4

•H

o

CD

4-1

T3

H

a

CO

a)

X!

rH

1

^Q

CU

C

c

5-»

o

O

1 i i

c

•H

O

CO

•

cu

iH

CO

o

cu

CD

a

>

CO

CO

cu

^TJ

iH

rH

c

cd

3

D

c*

rd

CM

CO

CO

4-1

0)

C

cfl

X

O

o

•H

>>

M

■u

r-\

a,

CO

cu

o

o

4J

>-i

CO

o

CU

6

•H

4->

•H

e

•H

X

CO

O

CD

'r-l

>

<

cu

•H

a

4->

iH

CO

cd

CO

u

4-)

cj

C

o

3

cu

•u

O

CO

X

CD

13

CO

>-t

C

PLi

cO

CO

cu

c

P£i>

o

-4

•H

|

iH

4-1

00

<

CO

c

00

•H

cu

c

Fj

S-i

•H

J_j

d

£

cu

00

o

■P

•H

X

CU

Pn

CO

-a

za

IV

46

Cove core facies tends to be highest in anorthite component. In general, no
pattern exists toward albite-rich plagioclases. Bulk composition of alkali
feldspar is probably about 0^- qr based on a few determinations.

Magnetite is often the most abundant opaque phase in both plutons.
It is compositionally restricted to titanium-poor varieties (generally
less than 1 wt. %) that show no trend from rim to core. Most of it is
found as subhedral grains in f erromagnesian minerals, closely associated
with ilmenite. Although much of the ilmenite may be secondary, subhedral
grains within biotite and in the matrix testify to its primary origin,
especially in rocks where it constitutes the only oxide phase. Regard-
less of pluton, ilmenite is manganif erous and ranges from 3 to 10 wt %
in MnO. As in the case of the magnetite, no consistent trend is present
between the two plutons and for each facies.

Geologic Interpretation

The Bottle Lake Complex consists of massive, late Paleozoic granitic
intrusives emplaced into the greenschist facies rocks of the Merrimack
synclinorium. Petrologic similarities within the Complex suggest that the
two main granitoid rocks were derived from geochemically equivalent source
materials. Both plutons are alike in their essential mineralogy and key
index accessory phases (primary sphene, allanite, zircon, apatite). This
lack of distinction in the petrography of the Passadumkeag and Whitney
Cove plutons is also expressed in their major element variations and abun-
dances by a general overlap in composition and trend. Complete overlap is
shown by the lead isotopic composition of feldspar and whole-rocks of both

Sl^8<%!£8?6;fS°W*k a5.6^?7)! and Hl'fHT (38.2-
39.2) is consistent with the interpretation that the source materials for
the Whitney Cove and Passadumkeag River plutons are similar (Ayuso and
others, 1980b). Preliminary^strontium isotopic determinations suggest
that both plutons have Sr/ Sr initial ratios of about 0.7053, and
further support the lead and strontium isotopic similarity of the sources
represented by the Bottle Lake Complex.

Source rocks of the Bottle Lake Complex must be dominated by a proto-
lith of igneous character (I-type of Chappell and White, 1974) as suggested
by field relations, accessory mineralogy, major and trace element compo-
sition, and relatively low initial ratios for lead and strontium. The
composite nature of the Complex, the circular or elliptical shape of the
plutons, development of an aureole by contact metamorphism, the range in
the lithological variety of granitic rocks, and the presence of mafic-rich
inclusions in the Complex are in agreement with the field relations out-
lined by Chappell and White (1974) characteristic of I-type sources. Ad-
ditionally, the absence of relict aluminous phases (garnet, aluminosilicates,
cordierite, or muscovite) and presence of sphene and hornblende are also
concordant with such a source. The abundance of magnetite suggests the
prevalence of high oxygen fugacities common in I-type materials. Compo-

47

sitionally, these plutons are also consistent with an I-type source, as
they are only slightly peraluminous, and exhibit regular and linear var-
iation in Harker diagrams for major and trace elements for a broadgrange
of silica contents. Finally, typical I-type sources have initial Sr/
Sr ratios of less than 0.7080, in good agreement with the preliminary
value of about 0.7053 for the Bottle Lake Complex.

Lead isotopic studies are inconsistent with an origin directly from
mantle sources, granulite rocks in the lower continental crust, or from
upper continental crust as represented by sediments exposed near the
Bottle Lake Complex (Ayuso and others, 1980b). However, a source con-
sisting mostly of volcaniclastic rocks is supported by the studies of
Ayuso and others (1980b) and Loiselle and Ayuso (1980).

The Bottle Lake Complex may represent. granitic magmas from a volcani-
clastic source, generated by progressively higher degrees of melting and
probably emplaced in the same order as their generation. Minimum melting
resulted in liquids belonging to the granite of Whitney Cove. Intrusion
of the rim facies was controlled by the northeast-trending pre-Late Paleo-
zoic faults in the area; this was followed by intrusion of the core facies
which probably represents remobilized material. After intrusion of the
Whitney Cove pluton, the fault was reactiviated and resulted in the cata-
clastic zone. The amphibolite rocks in this zone may represent a diabase
dike intruded within the fault prior to emplacement of the granite of
Passadumkeag River. As in the earlier melting episode, the quart z-monzon-
itic rocks intruded after the rim facies and they also represent remobi-
lized, refractory material.

Renewed interest in the origin of granitic rocks in eastern Maine is
exemplified by studies contrasting field relations (Ayuso and others,
1980a), and geochemical characteristics of the plutons intruding the
Merrimack synclinorium (Loiselle and Ayuso, 1980; Ayuso and others, 1980b).
Wones (1976, 1980) suggested that the granitic plutons of New England and
Sierra Nevada were significantly different, and that they probably repre-
sented different tectonic regimes. Recent work by Fyffe and others (1980)
on granitoids from New Brunswick further support the absence of typical
continent-ocean tectonic regimes. Geochemical study of the granitoids
intruding the Merrimack synclinorium in eastern Maine show geochemical
gradients unlike those in the Sierra Nevada. The major distinction in
granitoids from eastern Maine exists between plutons intruding different
tectonic blocks on either side of the Norumbega fault system. Each block
could have produced magmas from different processes that are now coinci-
dentally aligned as a result of fault motion. If fault displacement is
minimal, production of granitic rocks in the northern Appalachians may
indicate a process different from continent-ocean tectonic regimes during
the Mesozoic and Cenozoic eras.

48

References

Ayuso, R. A., 1979, The Late Paleozoic Bottle Lake Complex, Maine: abs.
with programs, Northeastern section Geol. Soc. America., v. 11,
p. 2.

Ayuso, R. A., in press., Bedrock geology of the Bottle Lake Complex: U.S.
Geological Survey, Misc. Inv.

Ayuso, R. A., Loiselle, M. , Scambos, T., and Wones, D. R. , 1980a, Comparison
of field relations of granitoids across the Merrimack synclinorium,
eastern Maine: abs., The Caledonides in the U.S.A., ed. , D. R. Wones,
I.G.C.P. Project 27: Caledonide Orogen, p. A-17.

Ayuso, R. A., Loiselle, M. , Sinha, A. K. , and Wones, D. R. , 1980b, Studies
on sources of granitoids in the Merrimack synclinouium, eastern Maine:
abstract, EOS, Trans. AGU, v. 61, p. 412.

Chappell, B. W. , and White, A. J. R. , 1974, Two contrasting granite types:
Pacif. Geol., v. 8, p. 173-174.

Cole, J. M. , 1961, Study of part of the Lucerne granite contact aureole in

eastern Penobscot County, Maine: M.Sc, unpub. , University of Virginia,
76 p.

Doyle, R. G. , Young, R. S. , and Wing, L. A., 1961, A detailed economic

investigation of aeromagnetic anomalies in eastern Penobscot County,
Maine: Spec. Econ. Studies Series No. 1, Maine Geol. Survey, Augusta,
Maine.

Fyffe, L. R. , Pajari, G. , Cherry, M. E., 1980, The plutonic rocks of New
Brunswick: abs., The Caledonides in the U.S.A., ed. D. R. Wones,
I.G.C.P. Project 27: Caledonide Orogen, p. A8.

Larrabee, D. M. , Spencer, C. W. , and Swift, D. J. P., 1965, Bedrock geology
of the Grand Lake area, Aroostook, Hancock, Penobscot, and Washington
counties, Maine: Contributions to General Geology, Geol. Surv.
Bull. 1201-E, 38 p.

Loiselle, M. C. , and Ayuso, R. A., 1980, Geochemical characteristics of
granitoids across the Merrimack synclinorium, eastern and central
Maine: The Caledonides in the U.S.A., ed. D. R. Wones, I.G.C.P.
Project 27: Caledonide Orogen, p. 117-121.

Ludman, A. , 1978a, Preliminary bedrock and brittle fracture map of the
Fredricton 2° quadrangle: Maine Geological Survey, Regional
Map Series, Open File Map 78-2.

49

Ludman, A. , 1978b, Stratigraphy and structures of Silurian and pre-
Silurian rocks in the Brookton-Princeton area, eastern Maine:
Guidebook for field trips in southeastern Maine and southwestern
New Brunswick, ed. A. Ludman, N.E.I.G.C. Guidebook, 70th Annual
Meeting, p. 145-161.

Olson, R. K. , 1972, Bedrock geology of the southwest one sixth of the

Saponac quadrangle, Penobscot and Hancock counties, Maine: M.Sc,
unpub. , University of Maine at Orono, 60 p.

Streickeisen, A. K. , 1973, Plutonic rocks. Classification and nomenclature
recommended by the I.U.G.S. Subcommission on the systematics of
igneous rocks: Geotimes, v. 18, p. 26-30.

Scambos, T. A. , 1980, The petrology and geochemistry of the Center Pond
pluton, Lincoln, Maine: M.Sc, unpub., Virginia Polytechnic
Institute and State University, 40 p.

Wones, D. R. , 1976, Phanerozoic plutons of New England: abstracts with
programs, Geol. Soc. America, v. 8, p. 303-304.

Wones, D. R. , 1979, Norumbega fault zone, Maine: Summaries of Technical
Reports, v. IX, National Earthquake Hazards Reduction Program,
Open File Report 80-6, p. 90.

Wones, D. R. , 1980, A comparison between granitic plutons of New England,
U.S.A. and the Sierra Nevada batholith, California: The Caledonides
in the U.S.A., ed. D. R. Wones, I.G.C.P. Project 27: Caledonide
Orogen, p. 123-130.

Wones, D. R. , and Dodge, F. C. W. , 1977, The stability of phlogopite in the
presence of quartz and diopside: Thermodynamics in Geology, ed.
D. G. Fraser, p. 229-246.

ROAD LOG

Assemble at 8:30 AM in the parking lot of the Lincoln Post Office,
just directly across from the First Bible Baptist Church. Most of the
trip will follow unimproved, two-lane dirt roads almost exclusively used
for hauling timber. Please remember that truckers expect unimpeded access
and the right-of-way in these narrow, shoulder-less roads. Because of
limited maneuverability and safe-parking in these roads, up to five vehicles
will be accomodated on the trip. Participants should bring a bag lunch and
expect two short hikes, mostly along trails.

Mileage

0 0 From Post office, go north on Fleming St.

0.1 0.1 Turn E (right) on Depot St.

50

0.2 0.1 Continue past the traffic light. Follow signs to U.S.
Route 6.

0.3 0.1 Turn E (left) at the stop sign at intersection with U.S.
Route 6, also known as Lee St.

1.0 0.7 Note the abundant, low-lying outcrops of the granite of
Center Pond.

1.9 0.9 More exposures of similar granitic rocks.

2.4 0.5 Same type of rocks are exposed here.

2.8 0.4 Continue east on U.S. Route 6 and note on the left side of
the road the sign for Schrite's Caribou Pond Cabins.

3.0 0.2 Outcrops on both sides of road belong to the Center Pond
pluton.

3.5 0.5 Same type of granitic rocks

4.6 1.1 STOP 1: CONTACT OF COUNTRY ROCKS WITH THE GRANITE OF

CENTER POND

This outcrop shows the prevalent mineralogy and mode of in-
trusion of the Center Pond pluton. Many of the field rela-
tions and petrographic features exemplified by this outcrop
are also common in the granite of Passadumkeag River. Both
plutons are hornblende-biotite rich, two feldspar granitoids,
exhibiting similar accessory mineralogy and textures. Granite-
country rock intrusive relationships between the Center Pond
pluton and the Vassalboro Fm. are typically transitional in
nature and are exemplified in this exposure. The transitional
character of the contacts is also common in the Passadumkeag
River granite, although knife-edged transitions and others
where numerous pegmatites and aplites crisscross the country
rock are also well represented. Take a moment to study
this outcrop on the opposite side of the road before approach-
ing the outcrop to inspect it more closely. Note that country
rock is clearly predominant on one end (eastern) of the out-
crop, and that it grades into a transitional or mixed zone on
the opposite end (western). On the eastern end of the outcrop,
the Vassalboro Fm. exhibits a typical pin-striped, shaly,
slightly rusty, and heavily-jointed appearance at the contact
with the granite of Center Pond. The purple color developed
in these rocks reflects the titaniferous biotite formed dur-
ing contact metamorphism. Immediately adjacent to the granite-
country rock contact, the Vassalboro Fm. consists of pin-striped
bands of differing mineralogy. Dark and light bands corre-
spond to biotite and calcite-rich zones, respectively, co-
existing with diopside, amphibole, epidote, and quartz.

51

Large blocks (3m) of country rock and granite are common
within the transition zone and typify the mode of intrusion
prevalent in both the Center Pond and Passadumkeag River
plutons. Inspection of individual blocks of metasedimentary
origin suggest that the pluton intruded in a lit-par-lit
manner, disaggregating blocks of various sizes. Aplitic
stringers intrude along the bedding planes. Several examples
of this relationship are present in this outcrop in elon-
gated blocks about 1 m in size, still exhibiting orientation
similar to that in the country rock. This rock shows ex-
treme deformation close to the cataclastic areas and this is
expressed by intense recrystallization of quartz, kinking
and smearing of biotite, and dislocation of feldspars. Most
biotite is red from hematite stain, and it is commonly rim-
med by finer-grained opaque minerals. In spite of the in-
tense rotation, grinding, and kinking in plagioclase, these
feldspars retain evidence of a large core rimmed by a clear,
optically continuous rim. Hornblende still shows subhedral
to anhedral outlines with abundant biotite inclusions. Sphene,
allanite, apatite, zircon, and opaque oxides are present else-
where in this pluton; however, they are not conspicuous in
this crop. This outcrop also shows cataclastic textures in-
terpreted by Scambos (1980) as part of a NE-trending fault
zone through the pluton with an offset of 2 km in a right
lateral manner. Note the strongly foliated character of the
rock (N80° E, 60-90 SE) as expressed by the feldspars and
mafic minerals on the western end of the exposure. This
fault zone may also be evident in the granite and country
rocks by the abundance of joints oriented generally to the
NE, and by a few thin (<5 cm) quartz veins.

Granitic rocks are generally gray in color, lack a porphy-
ritic or quench marginal zone at the contact with the country
rock, and preserve a nearly equidimensional character in their
fabric. While xenoliths of metasedimentary origin decrease in
the granitic rocks away from the transition zone, small clusters
(<10 cm) of mafic minerals (biotite + amphibole + sphene) and
plagioclase are common throughout the granite. These clusters
probably do not represent fully digested country rock as sug-
gested by their mineralogy, texture, and lack of alignment.
Because pegmatites and miarolitic cavities are absent and
aplitic rocks are uncommon in rocks near the contact, the
magma probably intruded without being saturated with volatile
constituents. The contact metamorphic effects are constrained
to within 1.5 km of the contact. In spite of the cataclastic
deformation in these rocks, it is still possible to ascertain
that at the time of intrusion, the liquid was probably satu-
rated with respect to all the major mineral phases. This re-
lationship is evident in the nature of the inclusions in
alkali feldspars (plagioclase, biotite, amphibole, quartz)
and the general subhedral nature of the phases to each
other in samples away from the fault zone.

6.2

0.6

6.5

0.3

7.1

0.6

7.5

0.4

7.8

0.3

52

Proceed E on Route 6.

5.6 1.0 Lee township line. Continue E on Route 6. Hornfels in the
contact aureole of the Bottle Lake Complex form prominent
hills, regardless of lithology and especially in the immed-
iate envelope around the pluton. Marshy land and streams
are commonly found on the slopes facing away from the in-
trusive, while the opposite slope commonly contains ex-
posures of the granite-country rock contact. As we proceed
to the east on U.S. Route 6, keep in mind that the row of
nearby prominent hills on the right side of the road con-
sists of the contact aureole of the Bottle Lake Complex.

Outcrops of Vassalboro Fm. on left side of road.

Intersection with South Rd. Continue E on Rt. 6 and notice
outcrops on right side of road.

More outcrops on left side of road.

Jefferson Ski Area Slope is directly ahead.

Continue E on Rt. 6 and through the flashing light at the
intersection of Rt. 168 and Rt. 6. Lee Academy and Lee
Post Office are on opposite sides of street to your left.

8.2 0.4 Outcrops of Vassalboro.

8.6 0.4 Maine Forest Service station to your right. Continue E on
Rt. 6.

10.3 1.7 Springfield town line. Continue E on Rt. 6.

10.9 0.6 Note the outcrops across from the AMOCO station.

13.9 3.0 Continue E on Rt. 6. Go by Springfield Post Office on your

right, and intersection of Rt. 169 and 170 with Rt. 6.

14.6 0.7 Turn S (right) on Dead River Co. road (Dead River Tree Farm).

We are now approaching the contact between the Bottle Lake
Complex and country rock, essentially at right angles. Note
the progressive increase in size and abundance of granitic
rock boulders in the next several miles.

16.2 1.6 Almanac Mountain is directly ahead of us. It is the most

prominent topographic feature in the vicinity and may be
identified by the firetower and relay antenna on its pinnacle,

16.5 0.3 Continue S and go by Springfield Cemetery on your right.

17 0.5 Outcrops of Pre-Silurian sulfidic schist on the right side

of road.

53

17.1 0.1 Continue S past the Intersection of road going to the east.

You may be able to see Duck Lake and Bottle Lake directly
ahead and on the left side of the road.

17.6 0.5 Lakeville town line. Turn left (E) at the intersection with

road going to Duck Lake.

18.5 0.9 Turn right (S) at fork on road going to Duck Lake.

19.2 0.7 Turn left (E) at fork with Getchell Mtn. gravel road. Note

the abandoned school at the intersection (Gowell School).

19.6 0.4 Cross Getchell Brook

20.2 0.6 STOP 2: RIM FACIES OF THE PLUTON OF PASSADUMKEAG RIVER.

Please turn around and try to park as safely and as far off
the road as possible. Mr. H. Snow owns this land and needs
free access to the road.

This stop exhibits pavements and low ridges forming
essentially continuous granitic outcrop to the top of Getchell
Mtn. (1041 ft.). Outcrops are generally fresh and very
extensive. The road ends within 1500 ft. from the granite-
country rock interface and about 4500 ft. from Getchell
gravel road. Because of its orientation at right angles to
the contact and abundance of fresh outcrop, this traverse
constitutes an outstanding opportunity to study the field
relations, fabrics, mineralogy and interaction of the
granitic and country rocks. Walk up the mountain to the
end of the road, and on the way down inspect the outcrops.

Granitic rocks on Getchell Mtn. belong to the rim facies of
Passadumkeag River. The incompletely developed rim facies
of the pluton may be subdivided into smaller domains on
mineralogical and textural grounds. Note in Figure 2 that
the rim facies is absent in a large area near the northern
contact. All Getchell Mtn. rocks are hornblende-bearing,
two-feldspar granites constrained roughly to the area in
contact with pre-Silurian (?) sulfidic schist. Schlieren,
different types of inclusions, mafic segregations, and
irregular felsic masses are common in this traverse. In
most cases, the long dimension in the mafic segregations,
schlieren, and mafic inclusions strike roughly N 50°E,
parallel to the granite-country rock contact. This preferred
alignment is also expressed by the long dimension of the
alkali feldspars along the traverse. All the granitic rocks
show a prominent and consistent NE-trending alignment. Note
the presence of seemingly irregular masses of granitic rocks
of markedly porphyritic nature surrounded by the more typical
rim facies rocks in many of the outcrops. In addition to

54

these felsic inclusions, aplitic dikes and pegmatitic pods
are also numerous and exhibit no difference with proximity
to the contact. No consistent orientation is expressed by
these dike rocks as they show significant variation in
thickness, and contrast with the uniform orientation of the
mafic concentrations. The latter attain lengths up to 1 m
and show cyclic mineralogic variations from mafic to felsic-
rich layers. Mafic inclusions show a great range in size,
shape, and mineralogy. They range from the small (5-10 cm),
ovoidal, fine-grained, and randomly oriented biotite and
plagioclase masses common throughout the pluton, to biotite-
rich aggregates, and finally to prominently porphyritic
quart z-diorite rocks. Note that in spite of the apparent
randomness in orientation of these inclusions, their long
axes are commonly oriented roughly to the northeast.
Pegmatitic zones up to several meters in thickness are also
exposed on the western slopes of the mountain and these con-
sist chiefly of feldspar (15-20 cm) , massive quartz
(white and rose), pseudohexagonal biotite books (5 cm)
and sulfide minerals.

Major textural changes are exhibited by granitic rocks
near the contact, but unfortunately these are not exposed
along the road. A zone of fine-grained and phenocryst-
rich rocks is present at the granite-country rock con-
tact; these rocks are associated with numerous aplitic
dikes and pegmatites that randomly crisscross the country
rock. Study of individual blocks of country rock reveals
that the magma also intruded by sending dikes in a lit-
par-lit fashion. Gradual changes in color, texture, and
mineralogy are evident with distance from the contact.
At the contact, the granite is strikingly porphyritic,
mafic-poor, and grayish in color. Toward the bottom of
the mountain it becomes seriate, richer in mafic minerals
and pinker. Amphibole is rare near the contact, but it
progressively increases in abundance, grain size and de-
velopment toward the interior.

Accessory and varietal minerals consist of allanite, sphene,
apatite, zircon and opaque oxides and they show a sympa-
thetic increase in abundance toward the interior of the
pluton. Intergrowths of biotite with either plagioclase or
hornblende are common, especially near the bottom of the
mountain. Interfaces where large, inclusion-rich biotite
flakes join plagioclase are characterized by a zone of fine-
grained albitic plagioclase, quartz, and Fe-Ti oxides.
Biotite and euhedral hornblende form intimately related
clots suggestive of coprecipitation; however, it is also
common to find subhedral biotite inclusions within horn-
blende. The probable sequence of crystallization near the

55

granite-country rock contact may be generalized as follows:
zircon, opaque oxides (magnetite), apatite, allanite and
sphene, followed by biotite, amphibole, plagioclase, quartz
and alkali-feldspar. Even most outcrops near the granite-
country rock contact show this order of crystallization,
except for the absence of hornblende. Myrmekite and graphic
textures, together with the formation of chlorite, seri-
cite, hematite, minor epidote, secondary sphene and opaque
oxides are also evident through the traverse.

Textural and mineralogic changes are less striking between
the rocks along the traverse compared to those near the
granite-country rock contact. Granitic rocks of the rim
facies progressively attain the mineralogy and texture
typical of the quartz-monzonitic core rocks. This type
of gradual change toward the core of the pluton is common;
however, in the region north of Upper Sysladobsis Lake the
transition from felsic rim to core is substantially more
abrupt.

Continue W on Getchell Road to Gowell School, retracing
route toward Dead River Co. Road.

21.2 1.0 Turn right (N) at the fork.

21.9 0.7 Turn left (W) at the fork.

22.8 0.9 Turn left (S) on Dead River Co. Road.

22.9 0.1 Continue straight, past the road to the firetower on

Almanac Mtn.

23.5 0.6 Almanac Mtn. is immediately to our right; note that directly

in front are Sysladobsis, Bottle, and Keg Lakes.

23.9 0.4 Contact between the Complex and country rock is approximately

here.

24.7 0.8 Continuous outcrops of the core facies of Passadumkeag River

pluton on both sides of the road. At this location the
rocks are similar to the granitic outcrops near the bottom
of Getchell Mtn.

25.1 0.4 Continue straight (S) on this road. Keg Lake is at your

left and Bottle and Sysladobsis Lakes are directly ahead.

25.4 0.3 At the fork, take dirt road on the right. The other road is

only .3 miles long and ends at Bottle Lake boat ramp.

26.9 1.5 This is a five-way intersection. Take a left on dirt road

going south that is a continuation of the Dead River Co.
Dump Road. Starting at this point, Sysladobsis Lake will
be on the right and we may be able to glance at the lake on
the way to the next two stops. This road is the main access
for the great number of cottages on the lake.

56

27.3 0.4 Pug Hole in Sysladobsis Lake is on the right.

29.2 1.9 This road transects a hill with abundant exposures of the

core facies of the Passadumkeag River pluton.

32.8 3.6 Next stop is about 1300' from the fork in the road. Park-
ing and turning around is difficult in this area, especially
after a downpour. All vehicles should be turned around and
parked on the same road that was used coming into the fork.
Occupants should then congregate at the fork for a short
hike. Proceed E along the dirt road on the northeast slope
of Porcupine Mtn. for approximately 1300' and note outcrops
on both sides of road. These are essentially continuous
for about 1000' to the east, and extend to the top of Por-
cupine Mtn.

STOP 3: RIM FACIES OF THE GRANITE OF WHITNEY COVE

This outcrop exemplifies the rim facies of the pluton of
Whitney Cove and it consists typically of pink granitic rocks
of low color index and seriate texture. Note the abundance
of aplites, pegmatites, quartz veins, and granophyre dikes
in this area. Also note the relative depletion and limited
range of lithologies in these mafic inclusions compared to
the previous stop. Felsic dikes are variable in attitude
and thickness, exhibiting a tendency to subdivide and criss-
cross within a single outcrop. Pegmatitic masses in this
area contain coarse-grained feldspars (5 cm) and quartz, and
form pods without a consistent orientation.

This exposure is mineralogically similar to most rocks
belonging to the rim facies and consists of microcline,
plagioclase, biotite, quartz, and accessory minerals.
Intergrowths of biotite and plagioclase are reminiscent
of the rocks on Getchell Mtn.; however, they are more
abundant in this area and the zone of interaction is sub-
stantially wider. Myrmekitic and micrographic textures
are ubiquitous as is the granulated appearance (mortar
texture) enveloping many of the alkali-feldspars. Allanite
is conspicuously euhedral, zoned, and very abundant in this
area; euhedral opaque oxides are more common than in the
rim rocks of the Passadumkeag River pluton and include
magnetite, ilmenite, and pyrite.

Return to cars.

Continue N on this road.

36.0 3.2 STOP 4: CORE FACIES OF PASSADUMKEAG RIVER PLUTON

57

Continuous outcrop extends to the top of the hill as pave-
ments, typically with smooth surfaces and covered by a
thick lichen carpet. Exposures are also abundant in the
opposite direction, close to the shores of nearby Sysla-
dobsis Lake. Note the abundance of mafic inclusions,
aplite dikes, and the marked increase in the color index
of this rock. In addition to the ovoidal, fine-grained,
biotite and plagioclase-rich clusters randomly dispersed
in the granitic matrix, note the fine-grained, porphyritic,
and hornblende inclusions. The latter are usually the
largest mafic inclusions in this facies, and generally show
a wide distribution of sizes. A preferred orientation to
the northwest is commonly expressed by their long dimensions,
although this contrasts with the easterly alignment of the
ovoidal clusters and the overall massive fabric in the rock.
Mafic-rich bands consisting of abundant biotite and horn-
blende and striking to the northeast are also abundant in
this area; notice that in contrast to the prophyritic inclu-
sions, these bands show an intermingling with granitic
material and indefinite boundaries. Mafic content of the
prophyritic inclusions varies extensively, sometines result-
ing in rocks of lower color index than the surrounding
granitic rock. However, note that large euhedral hornblende
and alkali feldspar are common to the granitic envelope and
porphyritic inclusions. Pods of felsic rocks showing aplitic
texture are easily distinguishable from the abundant inclu-
sions of high color index by their mineralogy and amorphuous
appearance. Few aplitic dikes are exposed in this facies
of the pluton; however, several dikes outcrop in this area,
varying in thickness, color, and orientation even over small
distances. Joints are generally well-developed and in some
cases they are strongly stained by hematite.

This outcrop also exemplifies the mineralogical and tex-
tural contrast between the core facies of the Passadumkeag
River and the Whitney Cove pluton. Amphibole is a common
phase in this stop, and it is commonly present as large,
euhedral, black prisms. Together with biotite, they impart
the characteristic high color index found throughout the
core of the Passadumkeag River pluton. Alkali feldspars
are prominently displayed, often mantled by plagioclase
and enclose all other essential minerals. Two or more
growth zones punctuated by oriented mineral inclusions
are commonly shown by the larger alkali feldspars. Plag-
ioclase contains abundant biotite and hornblende inclusions;
in most cases plagioclase grains are smaller than alkali
feldspar and form plagioclase-rich areas that alternate with
quartz-rich clusters.

58

As in the case of biotite, hornblende forms at least two
generations distinguishable by texture, size, and abundance.
Euhedral hornblende commonly forms large, inclusion-rich
(biotite, sphene, apatite, opaque oxides, zircon), randomly
distributed grains; however, note that smaller hornblende
is also abundant in the fine-grained plagioclase and biotite-
rich clusters. As in the Getchell Mtn. case, the relationship
between biotite and amphibole is of mutual intergrowths and
imply coprecipitation from the magma.

Return to cars.

Continue straight (N) toward the 5-way intersection.

40.2 4.2 Back at the 5-way intersection. Turn right on the same

dirt road used coming in and proceed toward paved road.

41.7 1.5 Turn N (left) at the paved Dead River Co. road.

44.4 2.7 Turn W (left) on gravel road leading to firetower on

Almanac Mtn.

46.0 1.6 STOP 5: CONTACT BETWEEN PASSADUMKEAG RIVER PLUTON AND

COUNTRY ROCK

Park in the small parking lot, find a comfortable place to
relax and eat lunch.

Extensive outcrops of the contact relations between country
rock and the Passadumkeag River pluton are exposed on the
southern slopes of Almanac Mtn. ; these exposures are essen-
tially continuous to the marked break of slope that envel-
opes the mountain. We will inspect an area about 100 m.
(S20°W) from the firetower. Please follow the trip leader
as we proceed toward these outcrops and stay with the group.
This stop serves as an example of the type of relationships
associated with the granite-country rock contact. As in the
case of the pluton of Center Pond, the intrusive contact
is expressed along a transition zone characterized by in-
termingling of country-rock and granitic blocks and lit-par-
lit disaggregation; however, small (<10 cm) pegmatites, pods
and aplitic stringers of variable thickness are also abundant,
and these randomly crisscross the country rock resulting
in a brecciated appearance.

Several types of granitic rocks are evident near the
granite-country rock contact in the vicinity of Almanac
Mtn. These range from strongly foliated, equidimensional,
medium-grained, mafic-poor rocks, to poorly foliated,
porphyritic, and more mafic. A common attribute of all
granitic rocks is, however, that foliation parallels the

59

intrusive contact. Mafic inclusions are relatively scarce,

as compared to the core facies, and they apparently consist

mostly of small (< 10 cm) ovoidal, plagioclase and biotite
clots.

Granitic rocks immediately adjacent to the contact are
notably poorer in mafic minerals compared to the core of
the pluton. Feldspars, biotite and quartz are often sub-
hedral even at the contact, and suggest that the granitic
magma was already saturated with respect to them at this
intrusive level. Note, however, that amphibole is rarely
present in these rocks, and that the small concentrations
of black minerals are dominated by commonly euhedral biotite,

Hornfelsic rocks are generally rusty on weathered surfaces
and purple in fresh cut. Many of the outcrops on Almanac
Mtn. show evidence for remobilization and plastic deforma-
tion probably related to intrusion of the Bottle Lake Com-
plex plutons.

47.6 1.6 Turn N (left) on Dead River road going toward intersection

with Rt. 6.

50.8 3.2 Turn E (right) on Rt. 6 going toward the town of Topsfield.

51.6 0.8 Carroll Townline, continue E on Rt. 6.

55.2 3.6 Bowers Mtn. is at your right and consists of both hornfels

and granitic rocks of the Passadumkeag River pluton. Horn-
fels are well exposed on the top and continue down to the
southeast limb of the mountain.

57.2 2.0 Pre-Silurian (?) Sulfidic schist outcrops.

57.9 0.7 Kosuth townline, continue E on Rt. 6.

62.2 4.3 Continue E on Rt. 6, past the dirt road going to Maine

Wilderness Canoe Basin campground on Pleasant Lake.

64.2 2.0 The lake directly ahead and on the left side of the road is

East Musquash Lake; East Musquash Mtn. is the highest moun-
tain in the vicinity and it is identified by the firetower
on the top.

64.8 0.6 Turn S (right) at the intersection with dirt road, sometimes

known as Amazon Road. Note the two signs at the entrance
identifying it as property of Georgia-Pacific Corp., and
also giving directions to the cottages on the shores of
Pleasant Lake.

60

68.7 3.9 Continue S on Amazon Rd. Orie Lake is at your left and also

notice that East Musquash Mtn. is in the background.

69.4 0.7 Pavements of country rock are almost continuously exposed

near the road, grading into granitic rocks of the rim facies
of the Whitney Cove pluton. Granite-country rock contact is
constrained within 200 m. Granitic rocks are foliated
parallel to the contact.

69.9 0.5 Continue S on Amazon Rd. ; trail to the left goes to West

Musquash Lake.

70.4 0.5 Pavement of rim facies of granite of Whitney Cove.

72.1 1.7 Continue straight on Amazon Rd. ; trail to the right ends at

Pleasant Lake.

73.2 1.1 STOP 6: CATACLASTICALLY DEFORMED RIM FACIES ROCKS OF THE

GRANITE OF WHITNEY COVE

This outcrop shows the cataclastically deformed granitic
rocks in the fault zone transecting the pluton of Whitney
Cove and continuously exposed in a belt approximately 1.5
km wide from Orie Lake to Junior Lake. The fault zone is
cut by the pluton of Passadumkeag River on its southwest
extension, and it may connect with the north-trending fault
through the country rocks mapped by Ludman (1978a). The
Norumbega fault system and the fault zones through the Bottle
Lake Complex and Center Pond plutons show the same northeast
trending orientation, and probably represent rejuvenated,
major tectonic lineaments in the region.

Note the markedly splintery, jointed, greenish appearance
of this outcrop as a result of fault motion after intru-
sion of the Whitney Cove pluton. Cataclastic outcrops
are commonly cut by numerous epidote and quartz-filled
dikes, striking generally in the northeast direction, but
commonly exhibiting substantial variation in attitude.
The fault zone in this area is characterized by granitic
rocks that show brittle deformation confined to narrow
zones. Stresses are apparently dissipated in the immediate
vicinity of the zones of maximum deformation, and are exem-
plified by spindle-shaped quartz grains within the epidote-
rich dikes. Right lateral motion along many of the east
trending joint sets is generally confined to less than 50
cm of displacement, and it is best displayed in the tran-
sected aplitic dikes. Even the mylonitic dikes are dis-
placed by a few centimeters by this set of joints. Al-
though the fault belt is delineated by a zone of abundant
cataclastic rocks in this area, this zone is also expressed
by ubiquitous epidote-f illed dikes throughout the Whitney
Cove pluton; this is especially applicable to the area
between this fault zone and the Norumbega fault system.

61

Note the prevasive and massive alteration recorded by these
rocks. Even those feldspars away from the cataclastic dikes
are crisscrossed by thin, epidote-rich veins, and show marked
alteration effects; biotite is completely chloritized and
forms random clusters in the matrix. Granitic textures are
obliterated near the epidote-filled veins and are characterized
by feldspars showing gradual disintegration, shearing,
rotation, and offset; note that the feldspars are im-
mersed within plastically deformed bands of recrystallized
quartz and biotite. Many of the exposures in this area are
characterized by aplitic dikes that are commonly lined by
silicified dikes dotted by fine-grained sulfide minerals,
up to 10 cm in thickness, and oriented to the northeast.
Foliation in granitic rocks is often well-developed and in
good agreement with the orientation of epidote-filled veins.

74.3 1.1 Cross Rainey Brook

75.0 0.7 Continue straight on Amazon Road; trail to the right goes

to a state campground along the south shore of Scraggly
Lake on Hasty Cove.

75.5 0.5 Note the abundance of cataclastically deformed granitic

boulders in this area.

75.9 0.4 Pavement of rim facies of Whitney Cove pluton rocks.

76.4 0.5 Continue straight at the intersection.
76.7 0.3 Continue straight at intersection.
77.3 0.6 Continue straight at intersection.

78.1 0.8 Pug Lake in Junior Bay is at your right.

79.3 1.2 Continue straight at the intersection; road to the right

goes to Whitney Cove in Grand Lake.

79.5 0.2 Exposures of the porphyritic core facies of the Whitney

Cove pluton.

80.0 0.5 STOP 7: PORPHYRITIC FACIES OF THE GRANITE OF WHITNEY COVE

PLUTON

Outcrops of this facies are best exposed along the shores
of Whitney Cove in Grand Lake. Contacts between the core
and rim facies of this pluton are rarely sharp, but rather
exhibit a gradual and progressive transition into the rim
facies where observed in the large outcrops near Pork
Barrel Lake.

62

Distinction between the rim and core facies of the pluton
depends on the development of strongly porphyritic and
fine-grained character of the core rocks. Mafic minerals
are predominantly biotite, and as in the rim facies, horn-
blende is absent; also in common with the rim facies, bio-
tite locally forms conspicuously euhedral, and larger
pseudohexagonal plates. Ovoidal clusters of fine-grained
biotite and plagioclase are generally small (< 5 cm) but
more abundant than in the rim facies. Metasedimentary and
quartz-dioritic inclusions are relatively rare.

Phenocryst mineralogy is dominated by the feldspars, but
quartz and biotite are also represented as subhedral grains;
alkali feldspar is clearly predominant in size, abundance
and development over inclusion-rich (biotite) plagioclase.
Alkali feldspar encloses all other phases and it is some-
times decorated by them along growth zones. In samples
where the porphyritic texture is best developed, plagio-
clase is nearly absent from the rock, and euhedral alkali
feldspar is abundant; in spite of their idiomorphic
nature, most alkali feldspars show minor development
of scalloped edges resulting from embayment by the fine-
grained matrix. Rapakivi texture is typically present
but it is never abundant.

Biotite is the first major phase to crystallize as evi-
denced by the abundant pseudohexagonal grains and common
inclusions within plagioclase; biotite rims of these larger
grains are often ragged and intermingled with feldspars.
A second generation of biotite is generally ragged in
appearance, fine-grained, and forms an interlocking net
around all other phases. Accessory and minor phases are
represented by allanite, apatite, sphene, opaque (magne-
tite and ilmenite) and sulfide minerals. Crystallization
of biotite, plagioclase, quartz and alkali feldspar fol-
lowed the accessory minerals, and this sequence probably
represents the generalized order of appearance. The only
common occurrence of muscovite in the core facies is in
wide (>50 cm) aplitic dikes that are numerous in the area
surrounding Whitney Cove.

Return to cars.

Continue on Amazon Road.

82.1 2.1 Outcrops of porphyritic granite.

83.0 0.9 Note Georgia Pacific sign.

83.1 0.1 Cross stream.

63

84.2 1.1 Granitic rocks of core facies.

84.7 0.5 Nice pavements of granitic rocks of Whitney Cove granite.

85.2 0.5 Extensive outcrops on the right, almost continuous.

85.7 0.5 STOP 8: RIM FACIES OF THE GRANITE OF WHITNEY COVE

Large exposures of this facies are common on most hilltops,
and these are exemplified by the massive outcrops exposed
on the Pineo Mountains, directly to the south and by the
outstanding exposures common on the cliffs and islands of
Grand Lake. In spite of the relatively abundant outcrop
coverage, most rocks are deeply weathered and difficult
to sample appropriately. This stop is close to the inter-
nal contact of the rim and core facies, as well as to the
unexposed eastern granite-country rock contact. Note how-
ever, that this rock retains an overall massive appearance
except where the long dimensions of feldspar suggest a pre-
ferred orientation to the north. Metasedimentary inclusions
are progressively more numerous in this rock with proximity
to the contact. However, in addition to these inclusions,
note those of quartz-dioritic composition and restitic tex-
ture exhibit markedly inhomogeneous distribution even at
the scale of the outcrop.

Although most rim rocks show remarkable textural homoge-
neity, a few pods of variable size are reminiscent of the
finer-grained core facies. Epitote-f illed fractures and
joints are common in many of these exposures and serve
as the principal fractures from which numerous secondary
fractures originate. Close to these fractured areas,
quartz has strongly undulatory extinction and both biotite
and plagioclase show kink bands. Note the thick (^60 cm)
gray, aplitic dike showing indistinct margins with the
granitoid and remarkable zonation from rim to core. Mafic
minerals, chiefly biotite, are concentrated at the margin
with the granite and are replaced by a felsic band followed
by a pegmatitic seam (felsic and mafic minerals) toward the
central portion.

Contrast the color, texture, and mineralogy of this outcrop
with the previous stops. These rocks are undisputably pink,
equidimensional in texture and consist of subhedral biotite
(apatite + zircon + magnetite + ilmenite) in discrete clusters
in the matrix, usually in close relationship with plagioclase.
Feldspar mineralogy is still overwhelmingly dominated by
alkali feldspar; to, mostly as subhedral and typically deeply
embayed grains. Plagioclase, biotite and quartz are commonly
found as inclusions in alkali feldspar, typically in random

64

orientations without crystallographic control. Rapakivi
textures are not abundant, but where present, they consist
of peculiarly skeletal and embayed alkali feldspars, en-
veloped by a wide rim of plagioclase; inclusions of biotite
and plagioclase are common in these alkali-feldspar cores.
Idiomorphic, gray quartz clearly embays and predates not
only alkali-feldspar but also some plagioclase; clusters
of quartz mosaics tend to form an essentially monomineralic
and interconnected net that intermingles with feldspar
clots. Plagioclase shows polysynthetic twinning and a
generally euhedral, homogeneous core enveloped by a pro-
gressively more albitic rim. Together with biotite and
Fe-Ti opaques, plagioclase exhibits peculiar intergrowths
and deep embayments indicative of a reaction relationship.

88.0 2.3 Cross stream and continue , straight.

88.5 0.5 Continue straight at intersection.

88.9 0.4 Pavements of granite of Wabassus Lake.

STOP 9: (OPTIONAL) WABASSUS LAKE GRANITE

This medium-grained granite is found only within the
Norumbega fault zone. It is finer-grained than either
the rocks of the Whitney Cove pluton to the north or the
Cranberry Lakes pluton to the south. It is more altered
than either of those plutons, and has fewer mafic minerals.
It does not correspond to either of the three plutons south
of the Norumbega fault zone (Lucerne, Lead Mountain, Cran-
berry Lakes) or to the Whitney Cove pluton north of the
fault. As a result, we have been unable to obtain a mean-
ingful estimate of movement on the Norumbega fault zone.

92.2 3.3 Turn E (left) at intersection with paved road and continue

to the intersection of U.S. Rt. 1.

93. A 1.2 Cross Big Musquash and continue E.

95.7 2.3 Continue E (straight). Road at your right goes to Peter

Dana Pt.

97.3 1.6 Turn N (take left) at intersection with Rt. 1 and continue

N to Presque Isle.

END OF TRIP

65
TRIPS A-2 and B-l

GEOLOGY AND PETROLOGY OF IGNEOUS BODIES WITHIN THE KATAHDIN PLUTON

R. Hon

Department of Geology and Geophysics

Boston College

Introduction

The Katahdin Pluton is a prominently placed plutonic body within a
discontinuous bimodal belt of plutonic rocks extending northeastward from
Rangeley, NW Maine, toward south of Presque Isle, NE Maine (Fig.l, Fig. 2,
see also Fig.l of Ayuso, Wones -this guidebook). This belt, named here as
the Greenville Plutonic Belt (after the town of Greenville approximately in
the middle of the belt) , consists of a series of NE-SW elongated plutons
ranging in size from a few km to 70 km and ranging in composition from
ultramafics (dunites of Moxie Pluton: Visher, 1960) to granites (IUGS class-
ification) with under- representation of intermediate compositions.

The Greenville Plutonic Belt is defined here as a separate geological
entity principally due to its unique tectono-petrogenetic characteristics.
First, the belt marks an abrupt NW termination of large-scale Acadian pluton-
ism within the Appalachian Orogen. Second, it parallels a belt of penecon-
temporaneous acidic volcanic rocks preserved in several synclinorial structures
immediately to the northwest (Piscataquis Volcanic Belt, Rankin, 1968 -Fig. 2).
Third, within the belt there are several large gabbroic bodies (Fig. 2) which
point to a mantle-crust coupling of igneous activity during the Acadian
orogenic event. Finally, the belt cuts across several metamorphic isograds,
from a subchlorite zone in the NE of Maine to the sillimanite zone in the
SW (Fig. 3). Since the exposures of sillimanite grade are a reflection of
later, postorogenic isostatic rebound due to continuing erosion, it is
possible, by assigning reasonable lithostatic pressures to each of the iso-
grads to make an estimate of the magnitude of the rebound. These calculations
indicate that the present surface of the Greenville Plutonic Belt resulted
by a differential regional "uplift" (tilt) of 5% grading linearly upward
from NE to SW. Alternatively, the surface exposures in the Rangeley region
represent conditions and environments which, immediately following the
Acadian orogenesis, existed at a depth 8 km greater than the conditions of
similar exposures at the Katahdin Pluton. This difference gradually dimin-
ishes to zero toward the Katahdin Pluton. The Katahdin Pluton is located
further to the NE where it is situated between the shallower quartz diorite
of the Moxie Pluton and the volcanics of the Traveler Rhyolite. The vol-
canics of Traveler Rhyolite are included in the upcoming general discussion
of the Katahdin Pluton. This is for reason that the Traveler Rhyolite,
as it will be shown later, is comagmatic with the Katahdin Granite.

The area of interest to this report was mapped in the early 60' s by
D. Rankin and A. Griscom (Ph.D. theses - Harvard University) and in the early
70's by the author (Ph.D. thesis - M.I.T.). D. Rankin mapped the area of
Traveler Rhyolite; A Griscom the area of Katahdin Pluton; and R. Hon studied
both units petrologically, mineralogically, and geochemically.

66

o

CO

a.
S

E

O

o
o-

|0

3
<

O
ID'

1)
CD

"\

\

\

5-1

0

•H

67

QJ

M

• A

0

4->

•

•H

rH

QJ

3

0)

a

O

PQ

•H

4-1

cO

3

o

a

rH
fa

•H
3

E

<D

O

4-»

• A

0)

• #t

rH

3

CO

4-1

4-1

rH

rH

0)

CO

rH

•H

fa

4-1

CD

qj

>

•H

• o

£

PQ

3

0)

3

CO

QJ

rH

cO

4-1

3

o

(1)

rH

r-l

3

•H

•H

V-i

•H

bO

cu

•

a

U

>

S

CO

■U

3

3

•H

4-1

rH

CJ

M-l

QJ

CO

T3

3

CD

H

O

QJ

•H

cu

92

PQ

o

U

3

CO

e

>

CO

o

O

•H

a

3

>

3

T3

•H

CO

O

tM

QJ

cO

QJ

a

•H

4-1

o

P

•H

CO

o

3

3

3

4-1

CfrH

CO

•

O

8

3

CO

fa

CD

rJ

>

rH

4-»

4-1

0)

•H

P-i

CO

o

•H

u

Q

3

CJ

•H

3

0)

1

O

0)

CO

<H

CO

.d

QJ

>

rH

•H

CO

U

4-1

r-l

QJ

rH

fa

S

O

o

fa

n

•H

>

1

1

1

1

1

i

3

i-H

CN

m

<!-

m

vO

a)

CD

u

• •

o

c

CU

N

•
CN

QJ

QJ

J-i

3

00

•H
fa

1

•- cn* n

in to

DDU

LU

Z
<

z

ID

CO

LU

§

LL
O

Q.
<

O

5

o

-j

o

LU

o

Q

LU
N

—I

<

LU

68

50 miles

INDEX MAP AND
METAMORPHIC ZONES
(GEOLOGIC MAP OF MAINE, 1967)

Figure 3. Met amorphic zones in Maine.

Legend: la- sub chlorite zone; 1- chlorite zone;

2- biotite zone;

3- garnet zone;

4- staurolite zone;

5- sillimanite zone;

6- sillimanite & orthoclase zone.

NOTE: The met amorphic zones in S. Maine are omitted for clarity

69

TRAVELER RHYOLITE - KATAHDIN PLUTON IGNEOUS COMPLEX

The Traveler Rhyollte - Katahdin Pluton igneous complex consists of a
3.2 km thick sequence of rhyolitic volcanics of the Traveler Rhyolite (TR)
Rankin, 1968; (Rankin, this guidebook) and a shallow, composite, predominantly
granitic Katahdin Pluton (KP) . Much of the area of KP (over 95%) is under-
lain by Katahdin Granite (KG) with the remaining area divided between a
stock of Horserace Quartz Diorite (HQD) and two smaller intrusions of
Debsconeag Granodiorite (DGD) . All three of these smaller bodies (HQD, DGD)
intrude Katahdin Granite along the NW-SE trending West Branch Penobscot River
fault system (Fig. 4) suggesting that the intrusions were tectonically controlled,

Petrological and geochemical evidence suggests that within the TR-KP
complex there existed two genetically independent magmatic episodes. The
slightly younger episode includes voluminous ash-flow depositions of TR and
subsequent crystallization of KG, where-as the slightly younger episode in-
cludes the HQD and DGD.

TRAVELER RHYOLITE - KATAHDIN GRANITE SERIES

Traveler Rhyolite

The Traveler Rhyolite (TR) is an almost rectangular (20 km x 12 km) ,
largely tectonically sunken block of topographically prominent Lower Devonian
volcanic rocks, occurring near the NE extension of the Katahdin Pluton. The
dominant rock type of TR is a welded ash-flow tuff with varying degrees of
compaction (up to 1:20) and variable phenocrystic content. Based on the
latter, in particular the presence or the absence of quartz, the volcanic
pile is divided by Rankin, 1968, into the basal Pogy Member (quartz present)
and the overlying Black Cat Member (quartz absent) . Other phenocrysts in-
clude plagioclase, ferroaugite, biotite, fayalitic olivine, and opaques.
By applying various geo thermometers and geobarometers, the phenocrystic
assemblage indicate that the phenocrysts of the Black Cat Member form an
equilibrium assemblage at T=800°C, P(H20)=1100 bars and f(02) near the fay-
alite-magnetite-quartz buffer curve. The temperature for the Pogy Member
is approximately 40 lower. The character of the contact between TR and
Katahdin Granite is clearly intrusive, documented by numerous apophysical
injections of Katahdin Granite into the volcanics and by a narrow zone (100 m)
of recrystallized metarhyolites.

Katahdin Granite

2
Katahdin Granite underlies over 95% of the 1350 km large Katahdin Pluton.

The pluton is of a subcircular shape, conspicuously elongated in the NE-SW

direction parallel with the direction of the above estimated 5% postintrusive

regional tilt. Thus the SW extension of the pluton represents a section which

was originally approximately 3.5 km deeper than the corresponding section at

70

GEOLOGIC MAP OF KATAHDIN PLUTON

Figure 4. Geologic Map of Katahdin Pluton.

HLP - Harrington Lake Porphyrg
DGD - Debsconeag Granodiorite
HQD - Horserace Quartz Diorite
KG - Katahdin Granite

71

the NE. This suggests that the three-dimensional form of the Katahdin Pluton
is a flat laccolith not yet fully unroofed along the contacts with the
Traveler Rhyolite. This is further supported by the observed contact dips
and the lack of gravity anomaly associated with the pluton (Kane, pers. comm.),
Size estimates of the Katahdin Pluton laccolith are approximately 40 km in
diameter and 5 km thick. The general shape of the laccolith is also suggested
by the spatial distribution of various textural facies of Katahdin Granite.
Fig. 5 illustrates the idealized distribution of these textural varieties
within the pluton.

The core of the pluton is formed by a massive, structureless medium-
grained biotite granite - the Doubletop facies (Fig. 5). This facies grades
upward into the Chimney facies through a development of granophyric inter-
growths which becomes progressively more and more pronounced and finer with
higher elevations, ultimately accounting for more than 60% of the rock. At
higher elevations, toward the summit of Mt . Katahdin, the granite becomes
also miarolitic, again progressively more prominent toward higher elevations
and concurrently changing character from filled vugs into open vugs (Summit
facies). Textural variations across the horizontal profile, toward the side
contacts are characterized by an evolution of subporphyritic and rapakivi
texture, the South Brother facies. This facies, near the side contacts
becomes aplogranitic with occasional "nests" containing abundant linings of
black to dark blue tourmaline (Wassataquoik facies) . The observed thickness
of the "outer zone" facies varies from zero to a minimum of 450 m, which
is best explained by local differences in the cooling rates.

Aplitic dikes occur throughout the pluton and their frequency seems to
be independent of elevation. The frequency of pegmatitic dikes is however
inversely proportional to the frequency of miarolitic cavities. The peg-
matites are absent at higher elevations but become more frequent toward the
core of the pluton.

Also of interest is the relative erosional resistance of the "outer
zone" granites. The Summit facies, by its extreme resistance against
weathering provides an effective anti-erosional capping at higher elevations,
topographically identified in the northern part of the pluton as an elevated
plateau contoured by sharp edges, deep valleys and glacial cirques. The
plateau, surrounding near its southern edge the highest point in Maine
(Mt. Katahdin: Baxter Peak - 5267 ft.), dips gently toward the NNE from 4600
ft. to 2800 ft. over a distance of 8.8 km. The calculated mean slope of the
plateau (6%) correlates closely with the estimated postorogenic tilt in this
region (5% - see Introduction) . It is believed that the inclination of the
plateau is not intrinsic to the pluton but rather a result of later isostatic
crustal re-equilibration. Wherever the anti-erosional "shield" of the outer
zone granites erodes away, the granites of the Doubletop facies are subject
to a relatively quick erosion due to their poorer resistance. This explains
the existence of steep hillsides around the exposures of the Summit facies
and the topographic lows in the remaining part of the pluton.

72

£3 KGCH

E§§ HGW

CZ3 KGSB

) KGO

HGCT

DISTRIBUTION-

FACIES OF KAT. GRANITE

TEXTURE

Core

Double top KGD

Granitic

Outer zone

Chimney
Summit

South
Brother

KGC
KGS

KGSB

Wassataquoik KGW

Granophyric

Granophyric and
Miarolitic

Subporphyritic
with Rapakivi
Overgrowth

Aplo granitic

Porphyritic dikes

Cathedral

KGCT

Porphyritic with
Aplitic Groundmass

Figure 5. Schematic distribution of textural facies of Katahdin
granite within the Katahdin Pluton Laccolith.

73

Irrespective of its textural variations, the Katahdin granite is a
homogeneous, massive, medium to fine-grained biotite granite of constant
chemistry and mineralogy (Fig. 6): alkali feldspar OR70:AB28.5: ANI.5 (34%),
zoned plagioclase AN34 to AN25 (26%), quartz (34%), biotite (6%), and
accessory apatite, allanite, zircon, tourmaline, opaques. There is no known
occurrence of muscovite, garnet, or sphene. Various geothermometers and
geobarometers indicate temperature of crystallization at T=710°C, P(T)=
P(H20)=1200 bars, and f(0o) near the fayalite-magnetite-quartz buffer
curve. The pressure estimate compares well with the estimated 1100 bars
for the phenocrystic assemblage of the Traveler Rhyolite, suggesting that
the KG represents a solidified magma chamber underneath the Traveler volcano

HORSERACE QUARTZ DIORITE - DEBSCONEAG GRANODIORITE SERIES

Horserace Quartz Diorite

The Horserace Quartz Diorite (HQD) forms a small elongated stock intruding
the Doubletop facies of Katahdin granite in the western part of the pluton
(Fig. 4). The stock is about 5.5 km long, and at its maximum 2.0 km wide
with outcrops on both sides of the West Branch Penobscot River. The HQD in-
trusion parallels a prominent fault system extending NW-SE for a minimum of
30 km from Chesuncook Lake to the center of the KP. On both sides of the
HQD stock, also parallel with the fault system, are numerous lamprophyre-like
dikes which are mineralogically and chemically identical with the HQD.
The ascent of the quartz dioritic magma was therefore tectonically controlled
and the fault remained active even after the intrusion solidified.

From aeromagnetic data, Allingham, 1960, deduced that the intrusion is
symmetrical with contacts dipping outward at about 55° tio a depth of at
least 3 km. Projecting the contacts upward, the upper contact can be estimated
to have been about 1.5 km above the present exposures. The contacts with the
surrounding KG are consistent with the model of a forceful tectonically con-
trolled intrusion. Brecciation of the granite and granitic inclusions within
the HQD are commonly observed near the contacts.

The HQD intrusion is concentrically zoned with a composition of biotite-
amphibole quartz diorite at the margins grading inward into amphibole- bearing
biotite granodiorite. Color index varies accordingly from about 25 to about
7 in the center. The principal mineral components (Fig. 6) are plagioclase
(andesine to oligoclase) , quartz, alkali feldspar, hornblende, biotite, and
accessoric apatite, sphene, opaques and fine primary "droplets" of sulphides
(probably pyrrhotite and chalcopyrite almost exclusively in the cores of
amphiboles) . Fig. 6 shows compositional variations plotted on the IUGS class-
ification triangle as well as variation diagrams suggesting that the principal
cause of differentiation involves simultaneous removal of plagioclase and
amphibole. Estimates of f(02) yield values above or near the Ni-NiO buffer
curve .

74

■ kGD
• DGD
A HOD
SN Sierra Nevada

( Bateman t963)

biotitt

10

70
60

10)

20
10

plagioclase

alkali feldspar

Figure 6. TOP: At top IUGS
classification triangle
for acidic plutonic
rocks with plotted
model analysis of rocks
from the Katahdin Pluton.

BOTTOM: Mineral
variation diagrams in
Horserace Quartz Diorite.

quartz

i i

T I- t • ■ > ' ■ T *

4 8 12

amphibole

16

.9

PL/

(ARPL)
.7

.2 .4 .6

AMPH/(BI*AMPH)

.8

75

Debsconeag Granodiorite

Debsconeag Granodiorite (DGD) is a medium grained amphibole- bearing
biotite granodiorite found principally in a 80 km^ large body in the
central part of the Katahdin Pluton (Fig. 4). Another occurrence of DGD
is a small stock (1.5 km2) located in the southern part of the KP. The
larger of the 2 stocks is intruded by dikes associated with the HQD.

DGD in hand specimen is similar to the Doubletop facies of the Katahdin
granite but with a noticeably higher content of biotite and ever present
traces of amphibole. Modal analyses are plotted in Fig. 6. Typically, DGD
contains alkali feldspar (21%), cyclically zoned plagioclases : AN35 to AN18
(37%), quartz (31%), biotite (9%), hornblende (1%), and accessory apatite,
zircon, allanite and opaques.

Major Element Geochemistry

Major element analyses of representative samples of each of the rock
types are plotted on an AFM (wt%) diagram and their normative compositions
are plotted on an AB-Q-OR diagram (Fig. 7). From these diagrams a clear
distinction exists between the more reducing TR-KG sequence and the more
oxidizing HQD-DGD series. AFM plots are compared to the Skaergaard trend
(reducing trend) and to the Lower California Batholith trend (oxidizing
trend) . The trend shown for Lower California Batholith is also typical
of calc-alkaline magmatism associated with destructive plate margins (high
alumina basalt-andesite-dacite association) . From major element geochemistry,
it is also deduced that the Traveler Rhyolite at its origin was at 8 kb total
pressure, 4 kb water pressure and in equilibrium with two feldspars and
quartz. The low oxygen fugacity of the TR magma was presumably maintained
through equilibrium with graphite, which is a common component of Appalachian
eugeosynclinal sediments.

Trace Element Geochemistry

Fig. 8 shows chondrite normalized REE abundances of all the rock types
within the Katahdin Pluton. The REE, as well as other trace elements,
clearly indicate that the two magmatic sequences are of different origin.
Modeling of trace element abundances show that the Katahdin Granite can
be derived from the Traveler Rhyolite magma if 15-25% phenocrysts are re-
moved from the TR magma by fractionation. Debsconeag Granodiorite is best
explained by mixing of approximately 40% Traveler magma and 60% Horserace
magma. Trace elements also suggest that the TR magma could have been pro-
duced by 20% partial melting of eugeosynclinal sediments. Trace element
and major element abundances of the Horserace Quartz Diorite are consistent
with abundances of calc-alkaline associations of continental margins. If
this inference is correct then the HQD magma may represent magma derived
along a subducting oceanic crust.

76

. TR-KG

O DGD
ii HLP
U HQD

■ WE9T BRANCH DIKE SWARM
O MAFIC DIKE, EAST SIDE OF
DOUBLETOP MTN

SK Skaergaard Trend
LCB Lower California
Batholith Trend

• DGD

▲ HLP

■ HQD

O Moxie Granite

A Lenngton Granite

G Echo Pond Granite

Figure 7. TOP

BOTTOM:

AFM ( wt %) plot of chemical analysis for rocks within
the Katahdin Pluton.

AB-Q-OR plot of normative compositions of Katahdin
Pluton.

77

200

SEQUENCE

200 -

La Co

r~

Nd

— i 1 r~

Sm Eu Gd

Ho

Yb Lu

Figure 8. Chondrite normalized REE abundances for rocks of Katahdin
Pluton. Sample 127 is from the dike associated with the HQD
intrusion. KR-Kineo Rhyolite.

78

1

30 km

Gabbro oV^

KGS

LovJ
U02)

Figure 9. Schematic diagram of origin of Traveler Rhyolite -
Katahdin Pluton Igneous Complex.

79

Origin of Traveler Rhyolite-Katahdin Pluton Igneous Complex

Fig. 9 schematically illustrates the possible time sequence for the
origin of the igneous complex. Mafic magma of mantle origin (possibly
derived along a subducting plate) raises the temperature naar the base
of the crust to initiate partial melting. The resulting melt of rhyolitic
composition is then intruded to shallow levels. From here it is rapidly
moved to the surface through violent ash-flow eruptions. The residual
melt fractionates and on crystallization forms the Katahdin Granite which
is in turn intruded by smaller intrusions of quartz diorite (HQD) and
granodiorite (DGD) .

Statement

Since the exact route of the trip is somewhat weather- dependent } the
trip log will be distributed at the first assembly place: Togue Pond
Campground (outside southern entrance to Baxter State Park) .

80
TRIPS A- 3 and B-2

ALPINE GLACIATION OF MT. KATAHDIN

D. W. Caldwell
Boston University, Boston

Introduction

Mt. Katahdin is the highest mountain in Maine (1601m) and, with
a local relief of about 1450 meters, is one of the largest mountains
east of the Rocky Mountains. The mountain is composed of quartz
monzonite (Katahdin granite informally) and is part of a large Devoni-
an batholith that intrudes Lower and Middle paleozoic sedimentary and
volcanic rocks. The relief of the mountain can be explained in part by
the greater amount of erosion of the sedimentary rocks compared with
that of the more durable plutonic rock. However, the Katahdin pluton
underlies much of the lowland south and east of the massif, so differ-
tial erosion can not completely explain the relief of the mountain
above the surrounding countryside.

Erratics found by Tarr (1900) and Antevs (1932) near the summit
of Mt. Katahdin and by Caldwell (1972) on other mountains in the
region support the view that Mt. Katahdin was covered by continental
ice sometime in the Pleistocene. There is no direct evidence that the
highest elevations were covered by Late Wisconsin ice, although recent
work by Davis (1978) suggests that they were.

Features Formed by Alpine Glaciation

Cirques

In the Mt. Katahdin area cirques occur only on those mountains
underlain by Katahdin quartz monzonite. Most of the cirques (6) were
formed on the massif which includes Mt. Katahdin and the other high
peaks which lie above timberline. The 3 largest cirques are on the east
side of the mountain and have headwall heights which range from 720 me-
ters to about 100 meters. These 3 great cirques have flat to concave
floors and steep headwalls composed largely of bedrock. Postglacial
rockfall and avalanche debris does mask the lower slopes of the cirque
headwalls and sidewalls. The total aspect of these cirques is one of
remarkable freshness, especially when compared with other cirques in
northeastern United States assumed to be occupied by glaciers in the
Late Wisconsin (Craft, 1979).

Aretes

The three east -facing cirques are separated by aretes. The most
typical arete is Hamlin Ridge, which separates North Basin ciraue from
Great Basin cirque (Figure 1) . The arete which separates Great Basin
cirque and South Basin cirque, Cathedral Ridge, has been shortened and

81

o
+->

a
u
c

o

fc-C x

£ o

• H 4-s
^4

O £

O -H

rH

rr-)

C O

•H tr.

c3

u

• CD

U-i

c

g

•H

>

■H

cn
o

to

eg

•H
r-i

o c

o +->

0"""d
•H rH
rH 0
X5 'H
O 4H

<D

rH

•H

82

lowered by glacial erosion and mass wasting. The most spectacular
serrate mountain crest is the Knife Edge but it may not technically
be an arete because there is no cirque on its South side. However,
the long narrow saw tooth ridge crest and the over 2000 foot (720m)
drop into South Basin more than make up for this deficiency.

Moraines

The aspect of alpine glaciation on Mt. Katahdin about which there
is the greatest controversy concerns moraines found within and down the
mountain from the three largest cirques. Tarr (1900), Antevs (1932)
and Caldwell (1965, 1972) identified moraines in each of these cirques.
In addition these authors believed the large Basin Ponds moraine was a
medial moraine formed between the combined alpine glaciers from the
three cirques and the still active tongue of the Laurent ide ice sheet.
The common interpretation was that the alpine glaciers which formed
these moraines were both contemporaneous with (at the Basin Ponds
Moraine) , and postdated (at the moraines within the cirques) , the Late
Wisconsin ice sheet. Davis (1978) believes the Laurentide ice sheet
covered Mt. Katahdin during the Late Wisconsin but that no alpine
glaciers postdated the ice sheet glaciation. Davis does not believe
that there are moraines within the large cirques and interprets the
Basin Ponds Moraine to be a lateral rather than a medial moraine.

References

Antevs, E. , 1932, Alpine Zone of Mt. Washington Range: Auburn, Maine,
Merrill $ Webber, 118 p.

Caldwell, D. W. , 1965, Pleistocene geology of Mt. Katahdin, p. 51-57,
in Caldwell, D. W. (Ed.), New England Intercollegiate Geological
Conference Guidebook for field trips in the Mt. Katahdin Region,
Maine, 58th Ann. Meeting.

, 1972, The geology of Baxter State Park and Mt. Katahdin: Maine

Geol. Survey Bull., No. 12, 57 p.

Craft, J. L. , 1979, Evidence of local glaciation Adirondack Mts. New
York: Friends of the Pleistocene, 42nd Ann. Meeting, 75 p.

Davis, P. T. , 1978, Quaternary glacial history of Mt. Katahdin, Maine:
Implications for vertical extent of Late Wisconsin Laurentide
ice: Geol. Soc. America, Abstracts with Programs, vol. 10,
no. 7, p. 386.

Tarr, R. S. , 1900, Glaciation of Mt. Ktaadn, Maine: Geol. Soc. America
Bull. , vol. 11, p. 433-448.

83

Road Log

Mileage

0 Togue Pond Camps. Starting time is 6:30 A.M. There is
a fee of $5.00 per car without State of Maine license to
enter Park so double up in cars as much as possible. We
might try to pay for permits the previous night to save
time.

0.3 Enter Baxter State Park. Road is built on large esker.

0.4 Bear right at fork toward Roaring Brook campsite.

1.0 Rat Pond on left. If day is clear, the caravan will stop
10 minutes to allow photographs of Mt. Katahdin skyline.

1.8 Rum Brook.

5.0 Windey Pitch. Sandy washed drift and sandy till in which
in 1966 there were sand-filled ice wedge structures. We
can stop here on return trip. This ridge has no bedrock
exposures and is believed to be an end moraine.

6.8 Avalanche Brook and Avalanche Field.

8.4 Roaring Brook Campsite. Bear left and park in designated
area.

Trail Log

0 Trail log from Roaring Brook Campsite. Approximate mileage
and hiking times (does not include time at stops) are
given.

1.5 Stop 1. Halfway Rock (half way between Roaring Brook and
(45 min.) Chimney Pond Campsites). The sandy drift exposed on trail

and on small ridge to the right of trail has noticeable
content of erratic fragments. Figure 1 shows location of
this and other stops which are identified by numbers on
the photograph.

1.9 Stop 2. The Basin Ponds are dammed by the Basin Ponds
(1 hr.) Moraine (Figure 1, locality 2). The moraine is composed

of about 99% Katahdin quartz monzonite and is inferred to
have been deposited by ice issuing from the three great
cirques to the west. The moraine has also been interpret-
ed as a lateral moraine of the Late Wisconsin Laurent ide
ice sheet.

2.3 North Basin cutoff.
(1 hr.
15 min.)

84

3.0 Stop 3. Blueberry Knoll and North Basin cirque. Blue-
(2 hrs.) berry Knoll is at the mouth of what is probably the best
preserved cirque in New England (Figure 1, locality 3).
From Blueberry Knoll the Basin Ponds Moraine and asso-
ciated humnocky topography to the east are clearly visible.

The evidence bearing upon the origin of Blueberry Knoll
and the age of the last Alpine glaciation in the North
Basin cirque will be reviewed. If we appear to be making
good time, we should hike into the cirque to view the
topography on the cirque floor at close range.

3.2 Hamlin Ridge trail. If the weather (and the leader) is
(2 hrs. still holding up, we will make the somewhat arduous

15 min.) climb up the Hamlin Ridge arete, in order to better view
the features under discussion.

3.9 Stop 4. Hamlin Ridge. In North Basin near the head of
(4 hrs.) the cirque, but separated from the headwall by a depres-
sion some 60 meters wide and 6 to 12 meters deep, is a
well defined accumulation of boulders. The feature
resembles a protalus rampart, formed by rockfall rolling
over permanent snow banks.

4.0 Stop 5. Hamlin Peak. Imposing and instructive views of
(4 hrs. North Basin, Hamlin Ridge, and Basin Ponds Moraine to the
15 min.) east, South. Basin, the Knife Edge, and the Table Land

toward the south. To the west, on the other side of the
large basin called The Klondike, is a range of mountains
with an unusual pattern in the forests. This pattern is
produced by bands of dead but mostly standing trees.
Although these bands may have something to do with the
wind, they cannot be properly called blowdown or windthrow,
as they commonly are.

4.3 Stop 6. Howe Peaks trail. View down the length of North
(4 hrs. Basin cirque.

30 min.)

4.5 Caribou Springs. Caribou were common on the mountain
during the 18th and 19th centuries but were all killed
off by 1900. In the 1960's there was an abortive attempt
to repatriate the Caribou when several drugged, pregnant
cows were lifted by helicopter to near this spot. After
they came to, none were ever seen again, as far as I know.

5.2 Stop 7. The Saddle. Here we will separate the men (who
(5 hrs. will go down the mountain) from the boys (who may want to
15 min.) climb to the summit). The summit is a hard mile from here,
our cars somewhat over 5 miles away. Those who accompany
the leader down the saddle slide be very careful of loose
rocks .

85

6.5 Chimney Pond. Spectacular view of the headwall of South
(6 hrs.) Basin beyond the rangers camp.

10.0 Roaring Brook and end of trip.
(7 hrs.
30 min.)

86

BEDROCK GEOLOGY OF THE SHIN POND-TRAVELER MOUNTAIN REGION

By Robert B. Neuman and Douglas W. Rankin

U.S. Geological Survey, Washington, D.C., 20560

and Res ton, Virginia 22092

Introduction

Low-grade metamorphism and the presence of fossils make the geology
of northeastern Maine a key to understanding the more metamorphosed rocks
of southern New England. The Shin Pond-Traveler Mountain region contains
many of the critical elements of that key.

The region affords the longest and most complete, stratigraphic
section in the State; it contains many fossil localities including an
unusual occurrence of well-preserved Early Devonian terrestrial plants
(see Trip B5) , and it contains one of the largest masses of felsic
volcanic rocks in the United States. Because facies change abruptly
within short distances and because many critical relations are exposed in
inaccessible places, the features to be seen along the routes of the trips
are only random and incomplete samples of the information upon which the
understanding of the geology of this area is based.

The geology of the quadrangles to be visited was mapped at a scale of
1:62,500. Neuman (1967) mapped the Shin Pond and Stacyville quadrangles
for the U.S. Geological Survey; he has been especially interested in
Ordovician stratigraphy and paleontology. Rankin (1961) mapped the
Traveler Mountain quadrangle, with special emphasis on the volcanic rocks
of Traveler Mountain, for a dissertation at Harvard University; he was
supported In part by the Maine Geological Survey. The Island Falls
quadrangle, east of the Shin Pond quadrangle, was mapped by E. B. Ekren and
F.C. Frischknecht (1967) of the U.S. Geological Survey, using
electromagnetic equipment to supplement surface observations. In addition
to these reports, more than a dozen papers by these geologists on one or
another aspect of their work have been published. Further, all but the
mountainous area is covered by aeromagnetic maps of the U.S. Geological
Survey.

We wish to acknowledge the essential role of Arthur J. Boucot, now of
Oregon State University, in identifying and interpreting the many Silurian
and Early Devonian brachiopods, and thus in establishing the relative ages
of many of the units mapped. Graptolites in considerably fewer numbers
were identified by W.B.N. Berry (University of California, Berkeley) whose
assistance we also gratefully acknowledge.

The 1966 New England Intercollegiate Geological Conference in the
Mount Katahdin Region (Caldwell, ed., 1966) included field trips led by
Neuman and Rankin through much of the same area discussed here. Although
the relevant sections of the 1966 guidebook are largely repeated here,
information derived from additional work by Neuman, Rankin, and others has
been added, and the itineraries of the trips have been modified to fit the
time limitations of the 1980 program.

87

Major tectonic features

The major structures of the region are the large anticline that
extends northeastward from the Stacyville quadrangle across the southern
half of the Shin Pond quadrangle (the southwestern end of the Weeksboro-
Lunksoos Lake anticline of Pavlides and others, 1964) and the
complementary synclines to the northwest and southeast (fig. 1). Lower
Cambrian(?) and Lower Ordovician rocks are exposed in the core of the
anticlinorium. On the northwest flank of the anticline are Upper
Ordovician rocks and a Silurian sequence of distinctive calcareous
sedimentary rocks, conglomerate, and volcanic rocks that are overlain by
Lower Devonian siltstone, sandstone, the volcanic rocks of Traveler
Mountain, and the overlying sedimentary rocks that were derived from
them. On the southeast flank of the anticline, by contrast, the Silurian
rocks are largely a monotonous assemblage of slate, siltstone, and fine-
grained sandstone, without volcanic rocks, and with little limestone or
conglomerate; in this region no sedimentary rocks of Devonian age have
been identified.

Most of the rocks are deformed and metamorphosed to the chlorite
grade of regional metamorphism. Metamorphism and deformation are least in
the Traveler Rhyolite and the overlying Trout Valley Formation, the latter
being remarkably little disturbed*

The rocks of the Lower Cambrian(?) Grand Pitch Formation are
intricately folded and faulted and are more deformed than those of
overlying formations. Argillaceous rocks of the Grand Pitch possess a
well-developed cleavage, and sandstones are commonly sheared. In many
places cleavage is folded; in some of the cleavage folds the earlier
cleavage is cut by a second one. Argillaceous rocks and lnterbedded
sandstone at the base of the overlying Ordovician Shin Brook Formation at
a few places are not as complexly deformed as those of the Grand Pitch.
At other places the lower part of the Shin Brook contains conglomerate
composed of fragments of slate and quartzite almost certainly derived from
the Grand Pitch.

Both the deformation contrast and the clasts of the conglomerate
indicate that a tectonic event separated the deposition of these
formations. Deformation contrasts between rocks that may be correlative
with the Grand Pitch and overlying Ordovician rocks have been described
elsewhere in the northern Appalachians (Cooke, 1955; Riordon, 1957;
Larrabee and others, 1965, p. E-8). Such a contrast through this large an
area suggests tectonic activity of regional extent at some time between
the Early Cambrian and the Early Ordovician; the term Penobscot
disturbance was coined for this event (Neuman and Rankin _in Caldwell
(ed.), 1966, p. 9).

The effect of the Taconic orogeny in this area is indicated by
several features. For example, the absence of Ordovician rocks beneath
the Silurian in most places along the northwest flank of the Weeksboro-
Lunksoos Lake anticline might be attributed to Taconic uplift and
erosion. This event may also be responsible for the apparent wedge-out of

88

Webster
K/ Lake

7

rt-

Trout
Brook

g

Billfieh
Pond

s

£. Branch
Penobscot R.

en

CT*

Shin
Pond

Figure 1.

Geologic map of Traveler Mountain, Shin Pond, and part of the
Island Falls quadrangles, and structure section approximately
parallel to route of excursions; generalized from 1:62 500
scale maps cited in text.

89

EXPLANATION

SEDIMENTARY AND EXTRUSIVE ROCKS

TNTRUSTVE ROCKS

SOUTHEAST OF ANTICMNE

>
a
u

I

U CO

I

4> t>

4> -r*
i-l O

g

^ §

41 -rt

J-6

NORTHWEST OF ANTICLINE

3— Dtvc
Trout Valley Formation
Dtv-3hale, si It stone, and
sandstone; Dtvc-conglomerate

Dk

Katahdin Quartz Monzonite
Dk-granoblastic rocks; Dkg-
granophyric rocks

Dtb

Dtp

Dg

Traveler Rhyolite
Dtb-Black Cat Member-highly compacted
welded tuff with 10$ phenocrysts of
plagioclase and auglte. Dtp-Pogy Member-
moderately compacted welded tuff with
15$ phenocrysts of quartz, plagloclase
and completely altered mafic minerals

Granophyre
May be equivalent to
Traveler Rhyolite

as

a

Dm

Matagamon Sandstone
Thick-bedded, fine to medium-
grained feldspathic sandstone

Ds

Seboamook Formation
Graded beds of fine-grained
sandstone and dark siltstone

DSv

Mafic volcanic rocks

Suv

ZE.

Sul

~* SB

}f

a a
ggg

s

Ss <-^ Ssc

Unnamed formations
Suv-mafic volcanic rocks, includ-
ing flows, tuff, and conglomerate;
Sul-calcareous siltstone, limestqne,
and conglomerate

Allsbury Formation
Ss-slate, siltstone, and minor
sandstone; Ssc-sandstone and
conglomerate with minor slate

SI

Unnamed formation *
Conglomerate, sandstone, and
non-calcareous siltstono

SOq

ss

a

Oc

Rockabema Quartz Dlorite

SS

a
eg

CO

Unnamed formation -
Conglomerate, sandstone, silt-
stone and basalt

CORE OF ANTICLINE

Och

Wassataquoik Chert
Chert with subordinate felsic
pyroclastic rocks

ss

a

Ov

Om

Mafic volcanic rocks (greenstone)

Mstadiabase

Oeb

Shin Brook Formation
Tuff, tuffaceous sandstone, and
conglomerate

-CgP

Grand Pitch Formation
Gray, green, and red slate, quartzite
and graywacke

a
£

J-

Contact

Fault; ball and bar on downthrown side
of those of dominantly vertical displace-
ment; arrows showing relative displacement
of those of dominantly horizontal movement

. . . Route of trips B3 and BA

* Stop of trip B3

• Stop of trip B4

90

Ordovician rocks at the southwestern end of this outcrop belt. The
contrasting facies of contemporaneously deposited Silurian rocks on
opposite sides of the Weeksboro-Lunksoos Lake anticline Indicate that an
ancestral form of this anticline developed during the laconic and remained
to separate the Silurian basins of deposition* Fragments of the Rockabema
Quartz Diorite in Lower Silurian conglomerate on the southeast flank of
the anticline were probably locally derived and indicate the minimum age
of that intrusive.

The Acadian orogeny was the last major deformation to affect the
area. Through most of the region, Acadian structure is characterized by
nearly vertical, well-developed slaty cleavage and shear surfaces; folds
are the dominant major structures, but there are significant contrasts in
the style of folding on opposite sides of the Weeksboro-Lunksoos Lake
anticline, and faults are important features in some places. On the
southeast flank of the anticline most beds as well as cleavage stand
nearly vertical; axes of minor folds and bedding-cleavage intersections
are generally vertical. By contrast, on the northwest flank, bedding over
wide areas dips moderately, and major as well as minor folds have moderate
plunges. Curiously, over a considerable area east of Traveler Mountain,
minor folds plunge northeast whereas major folds plunge southwest.

The age of the Acadian orogeny relative to the age of the Traveler
Rhyolite and Trout Valley Formation poses some difficult questions. The
Katahdin Quartz Monzonite lacks a tectonic fabric and clearly intrudes
folded Lower Devonian rocks; it is a post-orogenic pluton. The
structurally competent Matagamon Sandstone and Traveler Rhyolite are part
of these folds, but these competent rocks show the effects of deformation
less than the underlying rocks. The ash flow of the Pogy Member of the
Traveler Rhyolite overrode unconsolidated sediments of Matagamon as
evidenced by sandstone dikes in the basal Pogy flows and possible
channeling of the Matagamon by the ash flows (Rankin, this guidebook) .
Pebbles of felsite in the upper few meters of the Matagamon provide
further evidence that volcanism began before the end of the deposition of
the Matagamon Sandstone. On the other hand, both Hon (this guidebook) and
Rankin (this guidebook) conclude that the Katahdin Quartz Monzonite is a
subvolcanic pluton of the Traveler caldera. Thus, there can be no great
interval of time beween the deposition of the Matagamon Sandstone, folding
by the Acadian orogeny, and the intrusion of the post-orogenic Katahdin
pluton. These observations lend further support to the suggestion by
Naylor (1971) that the Acadian orogeny was a short-lived event.

The Trout Valley Formation is so little deformed that it may postdate
the Acadian orogeny. The age of the Trout Valley, relative to the
intrusion of the Katahdin Quartz Monzonite, is not known.

Stratigraphy

CORE OF ANTICLINE

Grand Pitch Formation (Neuman, 1962): Gray, green, and red slate and
siltstone and about equal amounts of vitreous quartzite and lesser amounts

91

of graywacke and tuff. Contains the trace fossil Oldhamia smithi
Ruedemann in red slate at several places along the East Branch of the
Penobscot River. Similar rocks of the Nassau Formation in New York that
have no fossils other than Oldhamia are considered to be of late
Proterozoic age (Fisher, 1977); Oldhamia also occurs with Early Cambrian
body fossils in the Weymouth Formation in Massachusetts (Howell, 1922),
and with late Early to early Middle Cambrian acritarchs in the Bray Group
of southeastern Ireland (Bruck, Potter, and Downie, 1974, p. 80). Minimum
thickness, 1500 m (5,000 ft).

Shin Brook Formation (Neuman, 1964): Tuff, tuffaceous sandstone and
conglomerate, breccia, and flows. Tuff, the most common rock, is massive,
greenish-gray, and porphyritic; contains saussuritized stubby to anhedral
plagioclase phenocrysts as much as 2 mm in cross section; it is of
intermediate composition, in the andesite-dacite range. Fossils, mostly
brachiopods, and fewer trilobites, bryozoans, gastropods, and sponges,
occur in the sandstone and tuff at different levels from place to place.
Paleontological studies of these and related fossils from New Brunswick,
Newfoundland, and Wales indicate a late Early Ordovician age (e.g.,
Neuman, 1964, 1976; Neuman and Bates, 1978; G.S. Nowlan, written commun.,
1979). Thickness variable, 100 to 750 m (300 to 2,500 ft).

Ordovician mafic volcanic rocks (greenstone): Largely massive, dark
greenish-gray, locally pillow lava and flow breccia. Petrology summarized
from Hynes (1976, p. 1216): Some rocks are highly porphyritic and contain
both feldspar and pyroxene phenocrysts in fine fluidal groundmass.
Alteration of some feldspars varies from core to margin suggesting that
the feldspars originally were zoned. Some pyroxene phenocrysts have good
oscillatory zoning. Pyroxene commonly less than 20 percent of mode.
These observations indicate that the rocks are probably meta-andesites.
Many coarse-grained rocks have ophitic texture; fine-grained vesicular
rocks that have almost 40 percent pyroxene were probably originally
basalts. The presence of both basalts and andesites is supported by bulk
chemistry which ranges from 50 to 60 wt percent silica. Thickness where
present, 300 to 750 m (1,000 to 2,500 ft).

Wassataquoik Chert: Chert with subordinate felsic and mafic pyroclastic
rocks. Thin-bedded, medium- to dark-gray, greenish-gray, and red chert;
tuff and tuff breccia interbedded locally. Siliceous shale interbeds
contain graptolites of the Climacograptus bicornis and Orthograptus
truncatus var. intermedins Zones, and conodonts and inarticulate
brachiopods. Estimated thickness, 100 to 450 m (300 to 1,500 ft).

NORTHWEST FLANK OF ANTICLINE

Ordovician conglomerate, sandstone, siltstone and basalt: Polymict
boulder to pebble conglomerate containing fragments of volcanic rocks,
slate, quartzite, and quartz pebbles, with interbedded siltstone,
ankeritic near the top; basalt at the base and near middle are dark
greenish-gray, very fine-grained, with phenocrysts of plagioclase,
pyroxene, and olivine; some pillow structures (Rankin, 1961, p. 41).
Exposed principally along the East Branch of the Penobscot River

92

(including Haskell Rock Pitch), presumably overlain by Lower Silurian
conglomerate; wedges out northeastward* Contains brachiopods, trilobites,
and corals of Late Ordovlcian (Ashgill) age* Maximum thickness about
1,200 m (3,500 ft), but wedges out abruptly.

Lower Silurian conglomerate, sandstone, and siltstone; Thick-bedded
polymict quart zose pebble conglomerate, micaceous sandstone, and gray and
red siltstone and slate; wedges out northeastward* Conglomerate contains
large thick-shelled brachiopods, such as Pentamerus sp » and Stricklandia
lens ultima Williams. As much as 250 m (800 ft) thick.

Upper Silurian calcareous siltstone, limestone, and conglomerate: Light-
gray calcareous siltstone and fine-grained sandstone containing thin beds
and lenses of silty limestone; includes some reefal limestone at Marble
Pond and elsewhere, and coarser grained sandstone and conglomerate in the
northeast corner of the Shin Pond quadrangle. Fossils, especially
brachiopods, corals, and stromatoporoids locally abundant. Some
assemblages dated as Early or Late Silurian (late Llandoverian or
Wenlockian) age; others are more certainly Late Silurian (Wenlock or early
Ludlow) age. Probable minimum thickness, 150 m (500 ft).

Upper Silurian mafic volcanic rocks (apparently a thick volcanic
equivalent of the calcareous siltstone sequence described above): Massive
metamorphosed mafic volcanic rocks including pyroclastics, interlayered
with green tuffaceous slate and siltstone, conglomerate with red and green
matrix, and muddy sandstone; also minor amounts of reefal limestone, some
containing basaltic clasts. Scattered fossils in green tuffaceous slate,
green matrix conglomerate, reefal limestone and debris derived therefrom;
some assemblages dated as Late Silurian (early Ludlovian), others dated no
more precisely than Silurian or Devonian. Some pre-Silurian rocks may be
included. Thickness a thousand meters or more (several thousand feet).

Devonian or Silurian mafic volcanic rocks: Tuff, breccia with scoriaceous
fragments, and probably some flows. Possibly the same as Upper Silurian
volcanic unit, but lacks fossils.

Seboomook Formation (Boucot, 1961, p. 169): Graded beds of fine-grained,
cross-bedded sandstone, dark-gray siltstone, slate, and a few thick beds
of fine-grained feldspathic sandstone like that of the Matagamon
Sandstone. One exposure of gray sandy siltstone at the base contains a
few Early Devonian brachiopods. Primarily a submarine- si ope and
prodeltadeposit according to Hall and Stanley (1973, p. 2101). Thickness
variable: 1,200 m (4,000 ft) on East Branch of the Penobscot River.

Matagamon Sandstone (Rankin, 1965): Thick-bedded, fine- to medium-grained
feldspathic sandstone and subordinate amounts of siltstone and slate like
that of the Seboomook. Sandstone commonly well laminated and crossbedded;
some displays scour- and- fill structure. Load casts of sandstone in
siltstone ("pseudonodules") rare. The Matagamon is a sandstone facies of
the Seboomook. Fossils scarce except in occasional shell beds where Early

93

Devonian (Becraft-Or iskany) brachiopods are abundant. Primarily delta- top
and delta-front deposits of a westerly Prograded delta (Hall and Stanley,
1973, p. 2101). Thickness, 1,200 to 1,500 m (4,000 to 5,000 ft).

Traveler Rhyolite (Rankin, 1968): Flinty aphanitic rhyolite that breaks
with a conchoidal fracture and contains 10 to 15 percent of small (1 to 3
mm) phenocrysts. Color ranges from light gray through various shades of
greenish, greenish gray and bluish gray to nearly black. In general, the
darkest rocks contain the least altered phenocrysts and the best preserved
primary textures. Largely welded ash-flow tuff, minor breccia, and rare
airfall tuff and sandstone shale. Columnar jointing characteristic of the
welded tuff. Younger than the Matagamon Sandstone of Becraft-Oriskany age
and older than the overlying Trout Valley Formation of late Early or
Middle Devonian age. Youngest stratigraphic unit intruded by the Katahdin
Quartz Monzonite.

Bottino and others (1966) determined a Rb-Sr whole-rock isochron of
360 m.y. ± 10 m.y. ( X Rb87=1.39xl0~11year"1) for the Traveler Rhyolite.
This is in conflict with its stratigraphic position below the Trout Valley
Formation and with the 395 m.y. age of the Katahdin Quartz Monzonite.

The Traveler Rhyolite is composed of two members:

Pogy Member - Lower member. Moderately compacted welded ash-flow

tuff containing about 15 percent phenocrysts of quartz, plagioclase

and completely altered mafic minerals.

Estimated thickness, 900 m (3,000 ft).

Black Cat Member - Upper member. Highly compacted welded ash-flow

tuff containing about 10 percent phenocrysts of plagioclase and

augite.

Estimated thickness, 2,300 m (7,500 ft).

Trout Valley Formation (Dorf and Rankin, 1962): Light blue-gray to black
shale, siltstone, sandstone, conglomerate, and minor amounts of sideritic
sandstone and black sideritic ironstone. A massive conglomerate lentil,
probably a deltaic deposit, is present at the base along South Branch
Ponds Brook — the route traversed by Field Trip B3. Although pebble and
granule conglomerate is scattered throughout, conglomerate lenses are less
common in the upper part; boulder and cobble conglomerate is largely
restricted to the basal conglomerate lentil. No rock fragments other than
felsite have been observed in the conglomerate.

Fossils include plants (in some places so abundant that the rock
resembles a low-grade coal), ostracodes, estherids( ?) , and eurypterid
scales. Well-preserved terrestrial plants dominated by Psilophyton
indicated a late Early Devonian (Ones quethawan- late Coblenzian) age to
Dorf and Rankin (1962), but possibly an early Middle Devonian age to
Schopf (1964, p. D49) and Andrews and others (1977, p. 283). The
relatively undeformed condition of the Trout Valley may be due to its
post-tectonic age if it proves to be equivalent to the post-Acadian Middle
Devonian Mapleton Sandstone of Aroostook County; on the other hand, it may
be due to its shielded tectonic position above the thick competent
Traveler Rhyolite. Exposed thickness about 450 m (1,500 ft).

94

SOUTHEAST FLANK OF ANTICLINE

Allsbury Formation (Ekren and Frischknecht, 1967): Sandstone
conglomerate, and minor slate - feldspathic sandstone, polymict pebble and
cobble conglomerate, and gray slate and siltstone. The coarser
conglomerate, containing cobbles of porphyritic quartz diorite, like the
Rockabema, greenstone, quartzite, and other rocks, occurs in the fault
slices of the southeastern flank of the anticline; at one place
interbedded sandstone yielded Early Silurian (late Llandoverian)
fossils. Estimated minimum thickness, 1,500 m (5,000 ft).
Slate, siltstone, and minor sandstone - medium- to dark-gray, greenish-
gray, and red slate and siltstone, and a few beds of fine- to medium-
grained sandstone. Monograptid graptolites rare, including late
Llandoverian to early Ludlovi'an forms. Estimated thickness, about 3,000 m
(10,000 ft).

INTRUSIVE ROCKS

Ordovician metadiabase; Gray and greenish-gray, fine- to coarse-grained
metadiabase forming massive ledges. Forms as a sill above Shin Brook
Formation.

Rockabema Quartz Diorite (Ekren and Frischknecht, 19 67): Fine- to coarse-
grained, gray to greenish-gray, sheared and altered porphyritic quartz
diorite and granodiorite, characterized by phenocrysts of quartz and
feldspar as much as half an inch in cross section. Potassic feldspar,
some slightly perthitic, constitutes as much as one-third of the
feldspar. Total feldspar somewhat more abundant than quartz. Chlorite
and epidote pseudomorphic after biotite; calcite abundant in patches and
veinlets. Locally contains abundant large xenoliths of greenstone and
quartzite.

Devonian granophyre: Light-gray granophyre containing about 5 percent
phenocrysts of quartz, plagioclase, and biotite. Plagioclase phenocrysts
commonly in rosettes 2 to 3 mm in diameter* Groundmass granophyric or
spherulitic

Katahdin Quartz Monzonite (Neuman, 19 67): Hypidiomorphic phase
(=granoblastic phase of fig. 1) is massive medium gray, medium grained,
and consists of two-thirds feldspar (about three-fifths perthite and two-
fifths zoned plagioclase), one-third quartz, and 5 to 10 percent
biotite. Where altered, potassic feldspar is pink, plagioclase is
greenish, and chlorite replaces biotite. The quartz monzonite is
porphyritic locally, and contains pink-weathering perthite phenocrysts 5
mm long in a groundmass somewhat finer grained elsewhere. The granophyric
phase is vuggy, pink, and contains phenocrysts of biotite. Vugs contain
epidote, tourmaline, quartz, and potassic feldspar. Border phase on the
east is fine grained and contains abundant fragments of thermally altered
and partially assimilated sedimentary rocks.

A preliminary Rb-Sr whole-rock isochron age of 395 m.y.
( A Rb =1.39x10 year ) for the Katahdin Quartz Monzonite was determined

95

Naylor and others (1974). This is a somewhat older age than the K-Ar ages
obtained for the Katahdin east of Ripogenus Dam (356 m.y.) and a diorite
stock at Nesowadnehunk Deadwater (361 m.y.) within the Katahdin batholith
(Faul and others, 1963). The 395 m.y. figure is close to several other K-
Ar ages reported by Faul and others (1963) for plutons in northern Maine
and to the K-Ar age of 390 m.y. of a small pegmatite, presumably from the
Katahdin batholith, in border-phase breccia east of Mt . Katahdin.

References

Andrews, H.N., Kasper, A.E. , Forbes, W.H. , Gensel, P.G. , and Chaloner,
W.G., 1977, Early Devonian flora of the Trout Valley Formation of
northern Maine: Review of Paleobotany and Palynology, v. 23, p. 255>
285.

Bottino, M.L., Schnetzler, C.C., and Fullagar, P.D. , 1965, Rb-Sr whole-
rock age of the Traveler and Kineo Rhyolites, Maine, and its bearing
on the duration of the Early Devonian [abs.]: Geological Society of
America Special Paper 87, p. 15.

Boucot, A. J., 1961, Stratigraphy of the Moose River synclinorium: U.S.
Geological Survey Bulletin 1111-E, p. 153-188.

Bruck, P.M., Potter, T.L., and Downie, Charles, 1974, The lower Paleozoic
stratigraphy of the northern part of the Leinster Massif: Royal
Irish Academy Proceedings, v. 74 B, no. 7, p. 75-84.

Caldwell, D.W. (ed.), 1966, Guidebook for field trips in the Mount

Katahdin region, Maine: New England Intercollegiate Geological
Conference, 58th, 61 p.

Cooke, H.C., 1955, An early Paleozoic orogeny in the Eastern Townships of
Quebec: Geological Association of Canada Proceedings, v. 7, p. 113-
121.

Dorf, Erling, and Rankin, D.W., 1962, Early Devonian plants from the

Traveler Mountain area, Maine: Journal of Paleontology, v. 3 6, p.
999-1004.

Ekren, E.B., and Frischknecht , F.C., 1967, Geological-geophysical

investigations of bedrock in the Island Falls quadrangle, Aroostook
and Penobscot Counties, Maine: U.S. Geological Survey Professional
Paper 52 7, 36 p.

Faul, Henry, Stern, T.W., Thomas, H.H. , and Elmore, P.L.D. , 1963, Ages of
intrusion and metamorphism in the northern Appalachians: American
Journal of Science, v. 261, p. 1-19.

Fisher, D.W., 1977, Correlation of the Hadrimian, Cambrian and Ordovician
rocks in New York State: New York State Museum Map and Chart Ser.
no. 25, 75 p.

96

Hall, B.A. , and Stanley, D.J., 1973, Levee-bounded submarine base of slope
channels in the Lower Devonian Seboomook Formation, northern Maine:
Geological Society of America Bulletin, v. 84, p. 2101-2110.

Howell, B.F., 1922, Oldhamia in the Lower Cambrian of Massachusetts
[abs.]: Geological Society of America Bulletin, v. 33, p. 214.

Hynes, Andrew, 1976, Magma tic affinity of Ordovician volcanic rocks in

northern Maine and their tectonic significance: American Journal of
Science, v. 276, p. 1208-1224.

Larrabee, D.M. , Spencer, C.W., and Swift, D.J. P., 1965, Bedrock geology of
the Grand Lake area, Aroostook, Hancock, Penobscot, and Washington
Counties, Maine: U.S. Geological Survey Bulletin 1201-E, 38 p.

Naylor, R.S., 1971, Acadian Orogeny: an abrupt and brief event: Science,
v. 172, p. 558-560.

Naylor, R. S. , Hon, Rudolf, and Fullagar, P.D. , 1974, Age of the Katahdin
batholith, Maine [abs.]: Geological Society of America Abstracts
with Programs, v. 6, no. 1, p. 59.

Neuman, R.B., 1962, The Grand Pitch Formation: new name for the Grand

Falls Formation (Cambrian?) in northeastern Maine: American Journal
of Science, v. 260, p. 7 94-797.

, 1964, Fossils in Ordovician tuffs, northeastern Maine, with a

section on the Trilobita by H.B. Whittington: U.S. Geological Survey
Bulletin 118 1-E, 38 p.

, 196 7, Bedrock geology of the Shin Pond and Stacyville

quadrangles, Penobscot County, Maine: U.S. Geological Survey
Professional Paper 524-1, p. 11-137.

, 1976 [1977], Early Ordovician (late Arenig) brachiopods from

Virgin Arm, New World Island, Newfoundland: Canada Geological Survey
Bulletin 261, p. 11-61.

Neuman, R.B., and Bates, D.E.B., 1978, Reassessment of Arenig and Llanvirn
age (Early Ordovician) brachiopods from Anglesey, northwest Wales:
Palaeontology, v. 21, p. 571-613.

Pavlides, Louis, Mencher, Ely, Naylor, R. S., and Boucot, A. J., 1964,

Outline of the stratigraphic, structural, and tectonic features of
northeast Maine: U.S. Geological Survey Professional Paper 501-C, p.
C28-C38.

Rankin, D.W., 1961, Bedrock geology of the Katahdin-Traveler area,
Maine: Harvard Univ. (Ph.D. thesis), 317 p.

, 1965, The Matagamon Sandstone, a new Devonian formation in

north-central Maine: U.S. Geological Survey Bulletin 1194-F, 9 p

97

1968, Volcanism related to tectonism In the Piscataquis

volcanic belt, an island arc of Early Devonian age in north-central
Maine, ±t± Zen, E-an, and others, eds«, Studies of Appalachian
geology, northern and maritime: New York, Interscience Publishers,
p. 355-3 69.

Riordon, P.H. , 1957, Evidence of a pre-Taconic orogeny in southeastern
Quebec: Geological Society of America Bulletin, v. 6 8, p. 389-394.

Schopf, J.M., 1964, Middle Devonian plant fossils from northern Maine:
U.S. Geological Survey Professional Paper 501-D, p. D43-D4 9.

98
TRIP B-3

THE TRAVELER RHYOLITE AND ITS DEVONIAN SETTING,
TRAVELER MOUNTAIN AREA, MAINE

Douglas W. Rankin
U.S. Geological Survey, Mail Stop 926-A
National Center, Reston, Virginia 22092

The stratigraphic and structural setting of the Traveler Rhyolite have
been outlined by Neuman and Rankin in this guidebook. Before describing the
field trip, I would like to discuss the Traveler Rhyolite in more detail.

The main body of the Traveler Rhyolite occupies at structurally
depressed, roughly quadrilateral area about 13 by 19 km on the northwest
limb of the Weeksboro-Lunksoos Lake anticline (see fig. 1 of Neuman and
Rankin, this guidebook). The aggregate thickness of the volcanic pile
within the depression is at least 3,200 m (10,500 ft). The main body of
rhyolite is bounded by high-angle faults to the north and west and is
intruded by the Katahdin Quartz Monzonite on the south. The depression is
thought to be an ancient caldera called the Traveler caldera (Rankin, 1968).

The Traveler Rhyolite is typically a dark-greenish-gray or bluish-gray
to nearly black aphanitic rock that breaks with a conchoidal fracture. A
whitish weathering rind is common. The rhyolite is porphyritic and
contains 3 to 30 percent (commonly 5 to 15 percent) small phenocrysts 1 to
3 ram across.

Many outcrops of the rhyolite offer little evidence about the type of
volcanic activity that produced the rock. Thin sections help decipher some
relevant features of these outcrops, but many rocks are too recrystallized
to classify. Ash-flow tuff (both welded and nonwelded), tuff breccia,
bedded air-fall (?) tuff and crystal tuff, lava, and tuffaceous sedimentary
rock have been identified. Enough evidence is preserved, however, to
establish that the great bulk of the Traveler Rhyolite consists of welded
ash-flow sheets. Little sedimentary rock and no soil horizons have been
identified within the volcanic pile, suggesting that the pile was erupted
in a relatively short time. The dominance of welded tuff argues for
subaerial eruptions. That the rhyolite welded ash- flow sheets overlie
sedimentary rocks containing marine fossils (Seboomook Formation and
Matagamon Sandstone) requires a change in the paleoenvironment from marine
to terrestrial (Rankin, I960).

Evidence that the bulk of the Traveler Rhyolite is welded ash-flow tuff
comes from both outcrop and thin section. The features preserved include
flattened pumice lumps, deformed (flattened) shards, columnar joints and
massive units without obvious bedding. No glass remains in these Devonian
rocks, and welding per se cannot be demonstrated, but flattening certainly
can.

Collapsed pumice lumps, giving rise to a crude planar structure, are
obvious in many outcrops but difficult to observe in others. Excellent

99

examples of rhyolite containing collapsed pumice lumps will be seen along
South Branch Ponds Brook, in loose pebbles on the shore of Lower South
Branch Pond, and, to a lesser extent, in the rhyolite of Horse Mountain.
Some difference in the average size of the pumice lumps is apparent from
one part of the volcanic pile to another, or even from outcrop to
outcrop. Detailed mapping eventually may demonstrate that individual flow
units are traceable, but so far, tracing of flow units has not been
attempted.

Columnar joints are obvious in many outcrops and cliff exposures.
They range in diameter from about 2 cm to 2 m; the larger ones are as much
as a few tens of meters long. Spectacular thin columns, averaging about 8
cm in diameter and a few meters in length, are exposed along Dry Brook.
Even smaller columns, some of which are bent, are found on the North Ridge
of North Traveler Mountain. We will not have time to visit either
locality, but we will see we 11- developed larger columns at Horse Mountain
and along South Branch Ponds Brook.

Some mountain slopes have a terraced appearance, suggesting the
erosion of a layered sequence. Commonly, the rises of these terraces are
zones of well-developed columnar joints, whereas the treads are zones of
indistinct columns or no columns. In a few sequences, the size of the
pumice lumps changes from one terrace rise to another suggesting that the
rises represent different flow units of a welded tuff sequence. More
commonly, flow units and cooling units cannot be distinguished from one
another or mapped (for terminology, see Smith, 1960).

The compaction foliation (eutaxitic texture) caused by the flattening
and welding of pumice lumps and shards is commonly roughly perpendicular to
the axes of the columns. Where this relationship holds true, the direction
of compaction was roughly perpendicular to the cooling surface, presumably
a quasi-horizontal surface, and the present attitude of the compaction
foliation is a measure of subsequent deformation. Columns having axes that
are not perpendicular to the compaction foliation may indicate that the
cooling surface was not horizontal or that the compaction foliation was
rotated by strain during deformation.

Tilted benches and measured attitudes of compaction foliation on Black
Cat Mountain, and on North Ridge, Center Ridge, and Pinnacle Ridge of
Traveler Mountain indicate that the axis of an open north- plunging
anticline runs through the South Branch Ponds. This fold is readily
apparent from the south end of Upper South Branch Pond on the trail up
Center Ridge. Compaction foliation dips of 20° to 40° are common.
Excellent columnar joints can be seen on the basal cliff of Center Ridge at
the northeast end of Upper South Branch Pond.

The Traveler Rhyolite is divided into the basal Pogy Member and the
overlying Black Cat Member. As explained below, the upper part of the
Black Cat Member probably could be mapped as a third member.

The Pogy Member is characterized by moderately compacted welded ash-
flow tuff containing about 15 percent phenocrysts, of which about one- third
are quartz and most of the rest are zoned plagioclase. Optical studies

100

indicate that the average composition of the plagioclase is about An/« and
zones range from An^c to Anrg (Rankin, 1961). Sanidine (about Or^y) occurs
with plagioclase and quartz in a few samples. Mafic silicate phenocrysts
may be absent or may constitute as much as 10 percent of the phenocryst
population. They are generally altered beyond recognition, but
clinopyroxene was observed in one thin section, and the remnants of biotite
phenocrysts were observed in a few thin sections. Garnet was seen in two
samples. Lithic fragments, including rhyolite, diabase, sandstone, and
shale, are present in most samples.

The Black Cat Member is characterized by highly compacted and welded
ash-flow tuff containing about 10 percent phenocrysts. Typically, 75
percent of these are zoned plagioclase (An«o to An,,, bulk composition
about Ab/y), 20 percent are augite (optically determined to be about Wo^o
Enog Fs#2)» and 5 percent are magnetite (Rankin, 1961). Quartz phenocrysts
are typically not present but rarely constitute as much as 10 percent of
the phenocrysts. Biotite phenocrysts coexist with augite in a number of
samples, and rarely, hornblende is the only mafic silicate present.
Fayalite coexisting with augite and biotite was observed in one thin
section, and garnet phenocrysts were seen in two samples. Growth
aggregates of phenocrysts, rare in the Pogy Member, are common in the Black
Cat Member (fig. IE).

An outlier of the Black Cat Member forms Soubunge Mountain about 14 km
southwest of Strickland Mountain. The preserved thickness of rhyolite at
Soubunge Mountain is about 250 m (800 ft), and it rests on a sandstone
similar to the Matagamon (Andrew Griscom, written commun., 1966). In some
of the Soubunge samples, hornblende phenocrysts are more abundant than
augite.

The ash-flow tuff on the summit of Big Peaked Mountain (Traveler
Mountain quadrangle), on Little Peaked Mountain, on Barrell Ridge, and
along Dry Brook is currently included in the Black Cat Member; it may be a
distinct unit at the top of the volcanic pile. In this tuff the pumice
lumps tend to be smaller and less compacted and to form a smaller
percentage of the rock than they do in the main part of the Black Cat
Member. Phenocrysts make up about 5 percent of the rock and are thus less
abundant than in the rest of the Black Cat. Plagioclase and augite are
present. The rock is highly jointed; some jointing is almost a fracture
cleavage.

The percentage of quartz phenocrysts is the only consistent difference
observable in the field between the Pogy and Black Cat Members. The
significance of the quartz phenocryst content was recognized after I
completed most of the fieldwork. Limited field checking has shown that the
subdivision is a valid one. It is, however, extremely difficult using a
hand lens to identify small quartz phenocrysts that constitute 5 percent or
less of the rock. For much of its length, the contact between the members
is approximated between locations from which hand specimens had been
collected previously.

From the evidence assembled, both the Pogy and Black Cat Members
consist domlnantly of welded ash-flow sheets. The phenocryst content of

101

the Pogy Is somewhat greater than that of the Black Cat and differs
significantly in that quartz is ubiquitous and sanidine may be present.
The Pogy Member is typically more altered than the Black Cat and more
commonly contains lithic clasts. The pumice lumps in the Black Cat are
characteristically more compacted than those of the Pogy. Length-to-
thickness ratios of 10:1 to 20:1 are typical for the collapsed pumice lumps
in the Black Cat. Ratios as high as 60:1 have been observed. In the Pogy
Member, on the other hand, ratios of 2:1 to 4:1 are more common.

As of this writing, the bulk composition of the Traveler Rhyolite has
not been well sampled. The three analyses published by E.S.C. Smith (1930
and 1933), the two analyses by Jun Ito, reported by Rankin (1961), and one
analysis by the U.S. Geological Survey (Rankin, unpublished data) do not
indicate significant difference in bulk chemistry between the Pogy and
Black Cat Members. This apparent uniformity in bulk chemistry is
consistent with trace-element data for several samples of the rhyolite
analyzed by Rudolph Hon (oral commun., 1975) as part of his Ph.D. thesis
study on the Katahdin batholith. In August 1979, I collected a suite of 19
samples, including 3 from the Pogy Member, 12 from the main body of the
Black Cat Member, and 4 from the upper part of the Black Cat Member. The
major-element analyses of these samples are not yet available.

The silica content of the available analyses ranges from 71.62 to
72.54 weight percent recalculated to 100 percent on a water-free basis) for
two samples of the Pogy Member and from 70.50 to 75.06 percent for four
samples of the Black Cat Member. The Differentiation Index of Thornton and
Tuttle (1960) ranges from 82.6 to 84.5 for the two Pogy samples and from
79.9 to 88.7 for the four Black Cat samples. To a first approximation, the
bulk compositions thus fall within the synthetic granite system. The
normative constituents for both members plot in the vicinity of the minimum
melting compositions in the steam-saturated system S102 - KAlSioOg -
NaAlSi30g (Tuttle and Bowen, 1958).

The model presented in 1968 (Rankin, 1968) to account for the observed
differences between the Pogy and Black Cat Members still seems
attractive. The arguments are summarized here from that paper. At the
time of eruption, both quartz and feldspar were crystallizing from the
magma that produced the Pogy Member, whereas liquids of the Black Cat
Member were crystallizing only feldspar. The slightly higher phenocryst
content of the Pogy further indicates that the Pogy was somewhat more
crystallized at the time of the eruption. These observations are
consistent with the hypothesis that the basal Pogy Member was erupted from
the cooler upper part of the magma chamber and that the overlying Black Cat
Member followed, probably relatively quickly, from deeper, hotter parts of
the magma chamber.

whether or not an ash flow will weld depends upon several variables, a
major one of which is emplacement temperature (Smith, 1960, and Boyd,
1961). Calculations by Boyd (1961) show that a magma having a lower
initial 1^0 content will erupt an ash flow having a higher emplacement
temperature than one erupted by a magma having a higher initial H^O
content. Other factors being equal, one would expect a drier magma to form
a more highly compacted welded tuff than a wetter one. In fact, the

102

emplacement and welding temperatures of the Black Cat Member were high
enough so that the tuff locally flowed after or during welding, as judged
by the rotated phenocrysts (fig. IF) and microscopic to mesoscopic flow
folds.

At equilibrium, the 1^0 content of a magma should increase upward in
the magma chamber (Kennedy, 1955). If the Pogy Member originated from
higher in the magma chamber than the Black Cat, it may have been wetter as
well as cooler. This suggested gradient in water content of the magma
chamber is also consistent with the phenocryst assemblage of the two
members. Tuttle and Bowen (1958) showed that at the liquidus , a decrease
in water content (decrease in steam saturation pressure) shifts the quartz-
feldspar field boundary toward the SICK corner of the SiCK - KAlSioOg -
NaAlSioOg system. That is, a liquid in equilibrium with quartz and
feldspar crystals at a higher steam saturation pressure might be in
equilibrium with only feldspar crystals at a lower steam saturation
pressure even though the composition of the liquid may be otherwise
unchanged. This relationship was incorrectly stated in the 1966 NEIGC
guidebook.

Corroborative evidence that the ELO content of the Pogy was higher
than that of the Black Cat comes from the widespread alteration of the
mafic silicate phenocrysts in the Pogy as well as the more commonly altered
plagioclase phenocrysts in the Pogy. This alteration may be deuteric,
caused by abundant volatiles rising through the Pogy ash flows after
emplacement. That the emplacement temperature of the Black Cat was higher
than that of the Pogy is indicated by the greater degree of flattening of
the Black Cat pumice lumps and the evidence for post emplacement flowage of
the Black Cat. The existing evidence is consistent with the Pogy Member
having had a lower magma temperature and emplacement temperature and a
higher I^O content (steam saturation pressure) than the Black Cat Member.
These suggestions can be related to gradients In temperature and i^O
content in the magma chamber.

Partial results from the analyses of the suite of 19 samples collected
in August 1979 indicate that the Traveler magma chamber was composition ally
zoned in components other than 1^0 . As stated above, the major- element
chemistry for these samples is not available at this time, the trace-
element analyses are only partially completed, and the Pogy Member is
poorly sampled. The work Is continuing. Quantitative analyses of trace
elements show that compositional gradients are most obvious in the main
body of the Black Cat Member but that more subtle gradients exist in the
pile as a whole. The results appear to confirm that the upper part of the
Black Cat Member on Big Peaked Mountain, Little Peaked Mountain, and
Barrell Ridge is a different unit. In summary, Sr and Zr (by x-ray
fluorescence), Nb (by spectrophotometry), and Li (by atomic absorption
spectroscopy) decrease upward through the volcanic pile, and U (by delayed
neutron activation analysis) increases upward in the pile.
Semiquantitative emission spectrographic analysis indicated that the boron
content of the Pogy Member is higher than that of the Black Cat Member.

Except for the outlier of the Black Cat Member on Soubunge Mountain,
nothing is preserved today of the Traveler Rhyolite outside the proposed

103

Traveler caldera. If we make a conservative estimate for the average
thickness of the rhyolite within the caldera of 1600 m, the volume of
rhyollte within the structural depression is about 400 km . A blanket of
welded tuff 250 m thick (the thickness on Soubunge Mountain) surrounding
the rim of the Traveler caldera to a distance as far as Soubunge Mountain
would more than double the volume of material erupted. If the original
volume of rhyolite was as much as 800 km , the Traveler rhyolite represents
one of the larger pyroclastic flow fields of the world (see Smith, 1960).
If the Katahdin pluton, which has an area of about 1,350 km (Griscom,
1976), represents the subvolcanic magma chamber as both Hon (this
guidebook) and I think, then 400-800 km of rhyolite is not unreasonable.

The Traveler Rhyolite is the northeasternmost and by far the largest
of a discontinuous belt of similar rhyolitic volcanic and shallow
granophyric Intrusive rocks of Early Devonian age that extends about 160 km
across north-central Maine. They are collectively called the Piscataquis
volcanic belt (Rankin, 1968). Five major volcanic centers have been
identified in the belt. Coeval mafic rocks, with the possible exception of
the Moxie pluton, are conspicuously absent from the Piscataquis volcanic
belt. Much of the rhyolite from these centers is peraluminous as evidenced
by normative corundum in the available major-element analyses and garnet
phenocrysts in the rocks. The Traveler Rhyolite appears to be the least
peraluminous (that is, to have the lowest percentage of normative corundum)
of the rhyolites, but garnet phenocrysts are present in a few samples of
the Traveler. Garnet phenocrysts from the rhyollte of the Piscataquis
volcanic belt range- in composition from about 80 to 90 percent almandine;
most of the remaining component is pyrope (Rankin, unpublished probe data).

I previously suggested that the Piscataquis volcanic belt was an
island arc (Rankin, 1968). The observations that the volcanic center's
overlie Lower Devonian marine sandstones, a shallower water fades than
that of the surrounding Seboomook turbidites, and that many of the
rhyolites must have been erupted subaerially are still valid. That is, the
Early Devonian rhyolitic volcanism took place along a welt or ridge within
a deeper water marine basin. That this volcanism was in any way related to
the subduction of oceanic crust now seems unlikely.

References

Boyd, F.R., 1961, Welded tuffs and flows in the rhyolite plateau of

Yellowstone Park, Wyoming: Geological Society of America Bulletin,
v. 72, p. 387-426.

Dorf , Erling, and Rankin, D.W., 1962, Early Devonian plants from the
Traveler Mountain area, Maine: Journal of Paleontology, v. 36,
p. 999-1004.

Griscom, Andrew, 1976, Bedrock geology of the Harrington Lake area,
Maine: Cambridge, Mass., Harvard Univ., Ph.D. thesis, 373 p.

104

Hitchcock, C.H., 1861, Geology of the Wild Lands, in Holmes, Ezekiel, and
Hitchcock, C.H., Preliminary report upon the natural history and
geology of the State of Maine: Maine Board of Agriculture, 6th Annual
Report, p. 377-442.

Kennedy, G.C., 1955, Some aspects of the role of water in rock melts, p.
489-504 _in_ Poldervaart, Arie, ed., Crust of the Earth: Geological
Society of America Special Paper 62, p. 489-504.

Rankin, D.W., 1960, Paleogeographic implications of deposits of hot ash
flows: International Geological Congress, 21st, Copenhagen 1960,
Report, pt. 12, p. 19-34.

,1961, Bedrock geology of the Katahdin-Traveler area, Maine:

Cambridge, Mass., Harvard Univ., Ph.D. thesis, 317 p.

,1965, The Matagamon Sandstone, a new Devonian Formation in north-
central Maine: U.S. Geological Survey Bulletin 1194-F, 9 p.

_, 1968, Volcanism related to tectonism in the Piscataquis volcanic

belt, an island arc of Early Devonian age in north-central Maine, Jin
Zen, E-an, and others, eds., Studies of Appalachian Geology, Northern
and Maritime: New York, Inter science Publishers, p. 355-36 9.

Smith, E. S.C., 1930, Contributions to the geology of Maine, Number IV. The
geology of the Katahdin area. Part 1, A new rhyolite from the State
of Maine: American Journal of Science, v. 19, p. 6-8.

, 1933, Contributions to the geology of Maine, Number 5. An

occurrence of garnet in rhyolite: American Journal of Science, v. 25,
p. 225-228.

Smith, R.L., 1960, Ash flows: Geological Society of America Bulletin,
v. 71, p. 795-842.

Thoreau, H.D. , 1950, The Maine Woods. Arranged with notes by D.C. Lunt:
New York, W.W. Norton and Co., 340 p.

Thornton, C.P., and Tuttle, O.F., 1960, Chemistry of igneous rocks. I.

Differentiation index: American Journal of Science, v. 258, p. 664-
684.

Tuttle, O.F., and Bowen, N.L., 1958, The origin of granite in the light of

experimental studies In the system NaAlSioOg-KalSioOg-Si02-H2^:
Geological Society of America Memoir 74, 153 p.

Itinerary, Trip B-3

Topographic quadrangle maps:

15 -minute 2-degree

Shin Pond Presque Isle

Traveler Mountain

105

Assemble at site of the former Shin Pond House, Maine Route 159, 9.5
miles northwest of Patten and 0.2 miles west of bridge over Shin Pond's
thoroughfare. The departure time is 8:30 a.m., Saturday, October 11,
1980. Be sure to attach yourself to the correct trip; more than one trip
will assemble here. Cars may be used on this trip but please be sure that
each vehicle carries at least four people, has a full tank of gas, and has
a useable spare tire fully inflated. Bring your own lunch. Note that
pets, including dogs, are not permitted in Baxter State Park. The trip
includes a 200-foot climb over a steep, forested scree slope and a 3-mile
trailless walk down South Branch Ponds Brook. Stout walking shoes are
essential; wet feet guaranteed for all but the most agile. Please stay
with the leader and in as compact a group as possible.

Mileage

0.0 Site of Shin Pond House, facing northwest along Grand Lake

Road. Trip B-4 also follows Grand Lake Road as far as the
East Branch of the Penobscot River. Some of Neuman's road
log for that trip as far as the Bowlin Pond Road is repeated
here. You should refer to that road log (this guidebook) for
more detail.

0.3 Roadside ledges are thin-bedded, crossbedded quartzite of the

Grand Pitch Formation and porphyritic Rockabema Quartz
Diorite.

0.4 Roadside ledges are medium- and dark-gray slate and quartzite

of the Grand Pitch Formation.

0.5 Very light colored and fine-grained phase of the Rockabema

Quartz Diorite and Grand Pitch slate.

1.5 T6R7 town line.

1.6 Road on right to Snowshoe and White Horse Lakes.

1.9 View of Sugarloaf Mountain straight ahead. The mountain is

capped by a metadiabase sill. The fossil if erous Shin Brook
Formation crops out on the slopes of the mountain beneath the
sill.

2.6 Crommet Spring lunch ground.

3.1 Ledges to left of road are of the fossiliferous Shin Brook

Formation.

3.6 Ledges on left are metadiabase sill that overlies the Shin
Brook Formation.

5.7 Ledges on left are quartzite of the Grand Pitch Formation.
5.9 Bridge over Seboeis River.

106

6.7 Road on right to Scraggly Lake.

7.1 Beginning of long straight stretch of road with view down

road of North Traveler and Bald Mountains, both held up by
the Traveler Rhyolite.
10.4 Side road right to Hay Lake.

10.7 Side road left to Bowlin Pond.

11.1 T5R8 town line. Gradational contact betwen Seboomook

Formation and Matagamon Sandstone crosses road near here.

11.8 Roadside ledges are of Matagamon Sandstone, as are all roadside
ledges as far as the shore of Grand Lake Matagamon.

13.1 STOP 1 . "Hurricane Deck." Overlook and exposures of

Matagamon Sandstone. The Matagamon Sandstone here is in a
northeast- trending structural basin, the Hay Mountain basin
(Rankin, 1965). These exposures are very nearly on the axis of
the basin, and the sandstone dips gently northeast. In good
weather, one can obtain a fine view here of the mountains to
the west and south. To the southwest, Mt . Katahdin is visible
between Turner Mountain on the left and Traveler Mountain on
the right. The long mass of Traveler Mountain is across the
valley of the East Branch. Although The Traveler (the highest
point of Traveler Mountain) is only 3,541 feet high, it rises
3,000 feet above the river. The bare conical peak of Bald
Mountain is set against North Traveler Mountain. The last
mountain to the right, barely visible from here, is Horse
Mountain on the shore of Grand Lake Matagamon. C.H. Hitchcock
(1861) referred to this as the mountain with the inelegant
name. Turner and Katahdin are composed of Katahdin Quartz
Monzonite, the others, of the Traveler Rhyolite.

14.7 Bridge over East Branch of the Penobscot River, a favorite

for white-water canoeists. H.D. Thoreau (1950) extolled the
joys of the East Branch after his 185 7 trip down it. The
store on the east side of the bridge was not there at the
time of my last pre- 19 7 9 visit, that is, at the time of the
1966 NEIGC trip. Civilization creepeth into the Maine
Woods. Road right on west side of bridge leads 0.5 mile
upriver to Grand Lake Dam at foot of Grand Lake Matagamon.
Good exposures of Matagamon Sandstone form east abutment of
dam.

15.7 Baxter State Park Boundary. This is the largest State park in

Maine and has an area of nearly 200,000 acres. 180-m cliffs of
rhyolite forming Horse Mountain on left.

15.9 STOP 2. Park as close to the edge of the road as you can.
Climb about 200 ft up steep slope to the base of cliffs. Be
extremely careful in crossing scree slope. Remember others are
behind you. The Pogy Member of the Traveler Rhyolite forms the

107

cliffs above. The Matagamon Sandstone underlies the scree
slope over which we climb. The contact, defined as the sharp
change in lithology from underlying obviously stratified rocks
to rhyolite above, is more or less exposed at the top of the
scree slope and dips 20° W. The top 6 m or so of the Matagamon
contain scattered pebbles of felsite and beds of tuffaceous
sandstone, indicating that some volcanic activity preceded the
main body of rhyolite.

The basal 0.6 or 1.2 m of the massive rhyolite are
composed of nonwelded tuff in which devi trifled shards are
clearly visible (fig. 1A). This nonwelded tuff grades up
into welded ash-flow tuff that appears to make up most of the
rhyolite of the Horse Mountain cliffs. Fragments of collapsed
pumice are visible in the rhyolite about a meter above the
base. Deformed and flattened shards are visible in a thin
section (fig. IB) collected from this locality 1.5 m above
the base. Columnar joints may be seen in the cliff face.
These are most obvious in the main part of the cliff to the
south and are perhaps most easily seen from the road. Some
are as much as 1.2 m in diameter and at least 12 m long.

If one traces the contact along the base of the cliffs, it
is seen to be an irregular surface having relief of as much as
5 or 6 m. This irregularity may be due to scouring by ash
flows, although in other places in the world, ash flows are
known to have crossed unconsolidated material without
disturbing that material.

16.3 STOP 3. Canoe landing on right opposite Maine Forest Service
Camp. Known locally as Eastern Landing. Walk 0.1 mile ahead
(north) along road. Roadcut in Pogy Member showing thin
anastomosing dikes of sandstone from the underlying Matagamon
Sandstone. Thicker clastic dikes have been found at the base
of the cliffs on Horse Mountain and on the shore of Grand
Lake Matagamon just ahead of us on the point. The largest
clastic dike is about 6 m thick and at least 10 m long

(as viewed from the bottom of the cliff). The clastic dikes
provide evidence that the ash flows overrode unconsolidated
sand of the Matagamon.

16.4 Gate house, Baxter State Park.

16.8 Road turns left away from lake and crosses ledges of rhyolite
of the Pogy Member.

17.9 Cross unexposed, high- angle fault between Traveler Rhyolite
and Seboomook Formation.

19.0 Trout Brook Farm, first cleared in 183 7, is now a Baxter

State Park campground. It produced hay for horses used in
logging operations until the 1940's. C.H. Hitchcock stayed
here in 1861. Rough side road, right, passes through farm

108

Figure 1. Photomicrographs of the Traveler Rhyolite. All are in plane-
polarized light. A., B., and C. are of the Pogy Member and are in
stratigraphic order from bottom to top. D., E., and F. are of the
Black Cat Member and do not represent a stratigraphic succession.

A. Nonwelded tuff, Pogy Member. Sample KT-153. Devitrified shards
showing axiolitic texture and unruptured "bubbles" are visible. Stop
2 of field trip; altitude of about 900 feet at base of cliff on Horse
Mountain about 0.5 m above the Matagamon Sandstone.

B. Welded ash- flow tuff, Pogy Member. Sample KR-73. Phenocrysts are
quartz and plagioclase. Flattened devitrified shards are visible;
some are deformed around phenocrysts. Clast of siltstone is in upper
right. Same locality as sample in figure 1-A about 1.5 m above the
Matagamon Sandstone.

C. Welded ash-flow tuff, Pogy Member. Sample KR-65. Texture suggests
vague outlines of highly compacted devitrified shards. The dark
mineral in the shards is probably celadonite. Altitude about 1,420
feet at top of cliff on Horse Mountain.

D. Welded ash-flow tuff, Black Cat Member. Sample Kt-499 . Phenocrysts
of zoned plagioclase and altered augite. Highly compacted and
devitrified pumice lumps impart the eutaxitic texture. The same
texture on a larger scale is visible in the outcrop. From second unit
above stream about 30 m downstream from Station 3 of traverse down
South Branch Pond Brook.

E. Welded ash- flow tuff, Black Cat Member. Sample KT-297. Highly
compacted. Phenocrysts of plagioclase, augite, and opaque minerals.
Growth aggregates of phenocrysts are visible. From ledges at an
altitude of 1,160 feet on east bank of stream flowing north from Black
Brook Mountains (Traveler Mountain quadrangle). This is about 2.8
miles west of South Branch Ponds campground.

F. Welded ash-flow tuff, Black Cat Member. Sample KT-208 . Phenocrysts
of zoned plagioclase, augite, biotite, and opaque minerals. Fayalite
is also present in the thin section. Extreme compaction of pumice
lumps. Rotated phenocrysts indicate that the material flowed after
compaction and welding so that texture resembles that of a lava. From
outcrop at an altitude of about 2,240 feet along an unnamed stream
that flows east from The Traveler toward Haskill Deadwater of the East
Branch.

109

•- %ir*

T"!

^«.

> '

~ ' 8

1 mm

1 mm

• -/<j»'

1 mm

Figure 1.

110

and continues to Webster Brook at the head of Grand Lake
Matagamon, crossing enroute some well-exposed open folds in
the Seboomook Formation.

19.5 Trout Brook on right parallel to road.

20.0 Sharp turn left. Ledges of Seboomook Formation in woods to
left. Excellent exposures of the Seboomook Formation just
upstream from the adjacent right-angle turn of Trout Brook.
Graded bedding, refracted cleavage, and many small folds
have been observed in these exposures.

2 0.2 Parking area left for trail to the delightful lakes of the

Deadwater Mountains.

20.5 STOP 4. Park along main road and walk 0.1 mile along side

road to site of old K. P. wooden dam on Trout Brook. The dam
is built on ledges of brecciated Black Cat Member, Traveler
Rhyolite. The high-angle fault bounding the rhyolite on the
north crosses the stream just below the dam. A thin wedge (9
to 12 m) of much fractured Matagamon Sandstone is north of
the fault. Beyond this is the Seboomook Formation.

22.9 Crossing of Dry Brook. Spectacular columnar jointing in the

upper distinctive part of the Black Cat Member about a mile
upstream. The columns are deformed into an open z-fold by a
normal fault of about 0.5-m displacement.

23.4 The Crossing. STOP. Leave appropriate number of cars so

that drivers may later be ferried to South Branch Ponds to
retrieve remaining cars. Considerable doubling up will be
necessary, but it is a short drive and no one wants to walk
both ways. Take lunches with you. After leaving some cars,
proceed up side road left to South Branch Ponds Campground.

2 5.6 STOP 5. South Branch Ponds Campground. Park cars in parking

area at entrance to campground and walk to shore of pond for
lunch. After lunch we will leave the cars here and walk down
South Branch Ponds Brook to The Crossing. There is no trail,
the distance is nearly 3 miles, and it is practically
impossible to make the trip with dry feet. Please stay with
the group. We must leave the campground by 1:30 p.m., and we
must all be at The Crossing no later than 4:00 p.m. The
sketch map (fig. 2) is traced from an aerial photograph, so the
scale is approximate. The log of the walk is by numbered
stations on the map (fig. 2), not distance.

South Branch Ponds Brook Wade.

1. Shore of Lower South Branch Pond.

South Branch Ponds occupy a glacial valley breaching a large
anticline in ash-flow sheets of the Black Cat Member of the Traveler
Rhyolite. On Black Cat Mountain to the west (right, looking up the

Ill

The Crossing

Base from (
aerial pho

9pr

^\17

16

V

L

\I4
h13

T

Vk/O

12 Jr

Y

11 (5
10 P^ \K»X*

6

\\o>

m

<#*■

^

ISi\

^OC

*»\

1000

2000 FEET

Approximate scale

South Branch
Ponds Camp
ground

Lower

South Branch

Pond

Figure 2. Sketch map of the South Branch of Ponds Brook area. Numbers
refer to localities described in the text.

112

lake away from the campground) , ash flows strike northeast and dip
moderately northwest. On Traveler Mountain to the east (left), ash
flows strike northwest and dip moderately northeast. The attitude of
these flows controls the northwest pattern of ridges on Traveler
Mountain. Neither the summit of The Traveler nor North Traveler is
visible from the lake shore. Mt. Katahdin is visible over the end of
the Upper Pond from the ridge north of the campground. Retrace route
out of campground past parking area and along road toward The
Crossing.

2. Reassemble on road at top of long hill (about 0.6 mile from shore of
lake). Trail right, which we will not take, leads to the Ledges, open
ledges of the Black Cat Member overlooking South Branch Ponds.
Columnar joints are distinct. Turn left downslope through open
woods. South Branch Ponds Brook is reached in about 0.2 mile. Turn
right and walk downstream.

3. Exposures of Black Cat Member of the Traveler Rhyolite. Note pattern
of concentric joints in outcrop on corner at stream level. Actually
there is more than one center about which joints are concentric,
giving rise to a pattern of intersecting curving joints. Distinct
columnar joints of small diameter are present above the stream on east
bank. Above this and slightly downstream is another flow unit having
columnar joints of larger diameter. Note the compaction foliation in
the unit (eutaxitic texture) brought out by the presence of very thin
lenticular bodies (collapsed pumice lumps).

4. First of a series of joint surfaces across which the stream flows.
Close examination will show that these are dip surfaces of ash
flows. Compaction foliation is roughly parallel to the surfaces, and,
in places, rather crude columnar joints are roughly perpendicular to
the surfaces. The ash flows strike about N. 70° E. and dip 30° N.
Local areas of crosscutting breccia are also present in this outcrop
on the left bank of stream.

5. Last of the series of dip slopes. Excellent swimming holes at bottom
of falls. Compaction foliation is visible on a number of steep joint
surfaces.

6. In stream bed on east bank just upstream from steep gravel bank at
corner. Lowest exposure of conglomerate of the Trout Valley
Formation. You, too, might call this a Pleistocene till upon first
enounter .

7. Jointing cuts cobbles in conglomerate. First clue that this is not a
Pleistocene till. Some cobbles are offset along the joints. Note
that clasts (pebbles, cobbles, and boulders) are well rounded and that
all of them are rhyolite. The clasts are so weathered that many of
them can be broken apart by hand. The weathering may date from the
Devonian Period. The conglomerate is crossbedded and dips gently
north, away from the volcanic rocks.

113

8. About 10 m of conglomerate exposed in the canyon wall. Where is the
contact with the overlying till? Note sandstone bed in the
conglomerate near top of exposure and lenses of black sandy
carbonaceous shale near bottom.

9. Junction with Gifford Brook.

10. Large exposure at curve of stream on left bank. Coarse conglomerate
no longer dominant. We are now above the basal conglomerate lentil
and in the main body of the Trout Valley Formation. Numerous black
chert lenses are visible and some have a vague internal structure.
Professor E.S. Barghoorn, of Harvard University, has identified one of
these as Prototaxites , which is generally regarded as of algal
affinities. Also present are siderite concretions and thin beds of
sideritic ironstone.

11. Upstream: sill of intermediate rock and a 15-cm-thick bed of
ironstone. Downstream: lens of carbonaceous black shale, from which
I first collected plant fossils in July 1955.

12. Light-gray-green fine-grained intermediate dike about 1 m thick.
Trends N. 30° W. a
sedimentary rocks.

Trends N. 30 W. and dips 40 N. Note chilled contact against the

13. This is "locality 4" of Dorf and Rankin (1962), from which the best
specimens of flattened spiny stems of Psilophyton were collected.

14. Gently dipping sill of intermediate rock about 3 m thick. Plant
remains can be found in nearly every outcrop of sedimentary rocks.
Reassemble here for half-mile walk out to Trout Brook. The group must
stay together from this point on. We will not follow the stream, but
will cut across country toward The Crossing.

16. "Locality 1" of Dorf and Rankin (1962). Long outcrop of gently
dipping, interbedded sandstone and shale of the Trout Valley
Formation. Sandstone is calcareous and current bedded. Rather well
preserved plant fossils have been recovered from some of the fine-
grained sandstone. Eurypterid scales also were found at this
outcrop. A fault of unknown magnitude cuts the southwest end of the
exposure.

17. The Crossing. Poorly preserved but large plant fossils occur in the
bridge abutment. Drivers will be ferried to South Branch Ponds to
recover cars.

End of trip. Return to Shin Pond and proceed to Presque Isle for
evening program.

114
TRIP B-4

ITINERARY

THE CORE OF THE WEEKSBORO-LUNKSOOS LAKE ANTICLINE, AND THE ORDOVICIAN,
SILURIAN, AND DEVONIAN ROCKS ON ITS NORTHWEST FLANK

Robert B. Neuman
U.S. Geological Survey, Washington, DC 20560

Topographic quadrangle maps:

15-minute 2-degree

Shin Pond Presque Isle

Traveler Mountain

Assemble at site of Shin Pond House, now burned, Shin Pond, Maine
ready for departure at 8:30 A.M., Saturday, October 11. Road limitations
dictate that heavy-duty vehicles be used. Please insure that (1) each
vehicle carries at least 4 people, (2) at the start the fuel tank of each
car is full, and (3) the spare tire is inflated and useable. Most stops
will be off the road, examining sections in the woods along streams. It
will be in the best interests of all participants to keep the group as
compact as possible. When the leader decides adequate time has elapsed for
examination of the exposures, he will return to the cars and proceed to the
next stop. Therefore, keep alert to the movements of the group, and do not
risk being left in the woods.

Mileage

0.0 Site of Shin Pond House, facing northwest.

0.3 Roadside ledges are thin-bedded, crossbedded quartzite of the

Grand Pitch Formation and fine-grained Rockabema Quartz
Diorite.

0.4 Roadside ledges are medium- and dark-gray slate and quartzite

of the Grand Pitch Formation.

0.5 Very light colored and fine-grained phase of the Rockabema

Quartz Diorite and Grand Pitch slate.

0.9 T5R7 town line.

1.3 T6R6 town line.

1.5 T6R7 town line.

1.6 Road on right to Snowshoe and Whitehorse Lakes.

1.9 View of Sugarloaf Mountain straight ahead. The mountain is

capped by a metadiabase sill in a syncline plunging
northeast. Beneath the sill, on the slopes of the mountain is
the Shin Brook Formation. The best fossils from the Shin
Brook were found on the easily climbed southern slope at 1,500
feet altitude.

115

2.6
2.7

Spring and lunchg rounds.

STOP 1. SHIN BROOK FORMATION (fig. 1 ): Stop cars near
Crommet Spring lunch ground; walk down road to left (west),
(recently relocated from allnement shown on sketch map) to
bridge over Shin Brook. Roadside exposures and east bank of
brook are Grand Pitch Formation: medium- to dark-gray slate
and thin-bedded, laminated, fine-grained light-gray quartzite,
some crossbedded. Note tight, steeply plunging folds,
especially in streambank.

Fossil locality

Fault, approximately located
U. upthrown aide: D, downtkrown side

'60
Strike and dip of bedding

Strike and dip of vertical beds

Beds face in direction of arrow

1000
I

2000
I

3000 FEET
I

Figure 1. Geologic sketch map of the vicinity of the type

section of the Shin Brook Formation; units 1-19 described in
text. Outlined areas indicate outcrops. From U.S. Geological
Survey Bulletin 118 1-E, fig. 2.

116

The base of the Shin Brook Formation is a spotted slate which
is about 60 cm (2 ft) thick. The type section measured here (taken
from U.S. Geol. Survey Bull. 1181-E, p. E-6) consists of:

Thickness
(Meters) (Feet)

Shin Brook Formation: 280 m (902 ft) measured

19. Tuffaceous sandstone and siltstone in graded
layers 8-30 cm (3-12 In.) thick, with
coarse-grained sandstone in the basal part
and finely laminated siltstone at the top;
siltstone more abundant than sandstone;
unit includes two layers of crystal tuff
16 cm (6 in.) and 60 cm (2 ft) thick,
respectively 10 35

18. Crystal tuff, greenish-gray; crystals are

green altered plagioclase 3 10

17. Covered 1 3

16. Tuffaceous sandstone, grit, and conglomerate;
finer grained part is well laminated,
coarser part not laminated and includes
fragments of porphyritic and nonporphyritic
fine-grained igneous rock; ledge in stream-
bed has distorted bedding structures that
suggest deformation prior to lithifi-
cation; base concealed 2 7

15. Tuffaceous sandstone, fine- to medium-
grained, calcareous; strongly sheared
with weathered pits that may have been
concentrations of fragmental fossils;
fossils largely brachiopods, too strongly
deformed to be identified 5 18

14. Covered 20 70

13. Crystal tuff, greenish-gray; with scattered
angular cognate rock fragments 1-15 cm
( li -6 in.) in average diameter; crystals
of both matrix and fragments are green
altered plagioclase; no primary layering
seen; quartz veins abundant 6 20

12. Covered . . , . - « 115 375

11. Crystal tuff; light-green altered plagio-
clase phenocrysts in a darker aphanitic
matrix; fractured 15 45

117

10. Covered 30 95

9. Tuff, fine-grained, light-greenish-gray;

abundant carbonate; strongly sheared, with

no bedding structures prerserved 10 30

8. Covered 10 30

7. Volcanic conglomerate, with granules of

aphanitic volcanic rock and dark slate,

strongly sheared; bedding obliterated 15 50

6. Tuffaceous sandstone, gray, medium- and fine-
grained; abundant carbonate; bedding
obliterated by strong shearing 6 20

5. Covered .»,---. , .,.*..*... 10 30

4. Volcanic granule conglomerate and coarse-
grained sandstone, light-gray, strongly
sheared 8 25

3* Sandstone and conglomerate, interbedded,

with conglomerate beds 25-50 cm (10-20 in.)

thick sandstone somewhat thinner; angular

to sub rounded fragments as much as 4 cm

(lV2in.) in average diameter include

volcanic rocks and fine-grained quartzite. . . .12 35

2. Phyllite, light-gray, probably tuffaceous 1 2

1. Slate, dark-gray, with small ( V4 - I2 ro™)

white grains (altered plagioclase?) with

rhombic outline abundant 1 2

Return to vehicles and proceed northwest on Grand Lake Road.

3.1 Ledges to left of road include fossil locality C of fig. 1

from which large specimens of Orthambonites robust us Neuman,
deformed Platystrophia sp., and bryozoans were collected.
Fossil locality D, shell beds of C_. robust us, is about 500
meters (1,500 ft) to the northeast in Crommet Brook.

3.6 Ledges on left are metadiabase of the sill that overlies the

Shin Brook Formation.

4.9 STOP 2. GRAND PITCH FORMATION AT SHIN FALLS: Walk south

along old road about 500 meters (2,500 ft) to second small
road to left (rushing water is plainly audible at this
point). Turn left (to east) and follow this road about 1,200
feet to Shin Brook at bridge (do not cross bridge); then
follow Shin Brook westward, downstream. The first large
ledges are greenish-gray, fine-grained quartzite with
interbedded gray and red slate. Cros shedding and graded

118

bedding are not as conspicuous as they are in the next
exposures downstream* Exposures at the upper cascades of the
falls have somewhat more slate and thinner quartzlte layers
than here. Graded bedding and crossbedding indicate that beds
face in the same direction for only short distances, and then
are abruptly reversed* Please do not descend waterfalls, but
return to road and cars from above log sluice*

5.7 Ledges on left are quartzlte of the Grand Pitch Formation.

5.9 Bridge over Seboeis River.

6.7 Scraggly Lake Road, TURN RIGHT.

STOP 3. SEBOOMOOK FORMATION AT SAWTELLE FALLS: Ledge on the
north bank of Sawtelle Brook is fine-grained sandstone
identical to that of the Matagamon Sandstone, one of several
that occur throughout the Seboomook. Walk eastward along old
road and trail about 2,000 feet to falls. The falls afford an
especially informative exposure of the Seboomook Formation as
large areas of bedding surface are visible* The ripplemarked
fine-grained sandstone is especially interesting, as it
reveals in plan view what is seen as small-scale crossbedding
at the base of graded sets in cross section* Note the regular
orientation of the ripplemarks and their elongation parallel
to the Intersection of bedding and cleavage. The plunge pool
and its downstream extension follow the trough of a syncline,
the beds on the opposite side of the stream dipping steeply
northwest. This is the only fold seen in the area that has a
horizontal axis.

Reverse direction; return to Grand Lake Road.

9.7 Grand Lake Road. TURN RIGHT.

10.6 Roadside ledge of Seboomook Formation.

11.0 Roadside ledge of Seboomook Formation.

11.3 Forest Service Camp, road right to Hay Lake; glimpse of lake

from the highway.

11.5 Bowlin Pond Road. TURN LEFT.

12.8 Weathered exposure of Seboomook Formation.

13.3 Roadbed "pavement" of Seboomook Formation.

13.4 Woods road left. Before parking for Stop 4 prepare to reverse
direction of caravan by proceeding past this road far enough
so that last vehicle can back into it, leaving sufficient
space so that all vehicles can be accommodated.

119

12.7 STOP 4. SILURIAN SEQUENCE AND CONTACT WITH GRAND PITCH

FORMATION. Follow woods road on which vehicles are parked
about 1,500 feet east and south to crossing of Bowlin Brook.
In roadbed at and near crossing note red shale (Grand Pitch)
interspersed with pebble conglomerate (Silurian), probably
tectonically. Proceed about 300 feet south down small gully
to its mouth at Bowlin Brook. Red and green slate and
siltstone of the Grand Pitch form the southeast wall of the
gully; southward along the brook for several hundred feet are
nearly continuous exposures of red, green, and gray slate and
greenish-gray, fine-grained, laminated, thin bedded quartzite
typical of parts of the Grand Pitch.

Debris on the gully floor, about 10 feet wide, covers the
trace of the contact between the Grand Pitch and the overlying
Silurian rocks, which are gray and dark-gray slate and slaty
siltstone containing thin seams of coarse-grained sandstone
exposed in waterfalls. Note the contrast of deformation on
opposite sides of the covered Interval; also note that the
lineation formed by the intersection of cleavage with bedding
surfaces of the Silurian rocks plunges to the northeast.
(Although the contact between the Grand Pitch and Silurian
rocks here has previously been considered to be an
unconformity, perhaps slightly faulted due to contrasts of
structural competency of the rocks juxtaposed here, geologic
relations in this vicinity, and the variety of units northwest
of the Grand Pitch along this strike belt, some almost
certainly overthrust by the Grand Pitch, suggest the
possibility of major fault movement at this level, perhaps a
decollement .)

About 50 feet west of the brook are beds of crudely
graded quartzose pebble conglomerate and coarse-grained
sandstone that are probably covered beneath the waterfalls in
the brook. Above the waterfalls are about 30 feet of red
slate and siltstone which are in turn overlain by about 50
feet of interbedded conglomerate, sandstone, and siltstone;
these conglomerates contain Pentamerus and other Silurian
brachiopods.

After examining this section, return to the Bowlin Pond
Road for examination of the remainder of the section.

The following section (revised from U.S. Geol. Survey
Prof. Paper 524-1) was measured in the woods west of the road
(fig. 2).

Thickness
(Meters) (Feet)
Seboomook Formation, 51 m (170 ft) measured,
of which 28 m (95 ft) is exposed):

12. Siltstone, dark-gray, in beds 10-15 cm
(4-6 in.) thick; basal parts formed

120

Seboomook
Formation

EXPLANATION

Contact

Shown on geologic map (pi. l) ; solid line
where exposed; long dashed where approxi-
mately located; short dashed where inferred

60

Strike and dip of beds

75

Vertical Inclined

Strike and dip of cleavage

Area of exposed bedrock

7427-SD
x

Fossil locality

N

Base of Silurian rocks,
approximately located

100 0 100 200 300 400 FEET
I I I I I I I

Figure 2 . Location of exposures and units of Silurian sequence

along Bowlin Pond Road, Shin Pond quadrangle, described in
text. Sc- conglomerate, Scs-lower calcareous siltstone, Sl-
limestone, Ss-upper calcareous siltstone. From U. S.
Geological Survey Professional Paper 524-1, figure 2.

121

of fine-grained crossbedded sandstone

1-3 cm (V2- lV2in.) thick 20 70

Covered interval 8 25

11. Silts tone, gray, sandy, noncalcareous;
bedding obscure but shown in subtle
contrast between more and less sandy
parts; scattered small brachiopods,
Metaplasia? sp. and Plectodonta sp 8 2 5

Covered interval 15 50

Upper Silurian Rocks, at least in part, 2 73 m
(820 ft) measured, of which 153 m (515 ft) is
exposed):

Covered interval 55 180

10. Silt stone, gray, calcareous; bedding surfaces
even, planar, at 5-25 cm (2-10 in.)
intervals; widely scattered comminuted
organic debris 25 85

9. Calcarenite, largely pelmatozoan

debris; silty partings irregular and

discontinuous; stromatoporoids and

favositid corals as much as 15 cm (6 in.)

in cross section 15 50

8. Siltstone, calcareous; contains silty lime-
stone nodules a few centimeters (inches)
in diameter, irregularly distributed 8 25

Covered Interval 20 70

7. Siltstone, calcareous, gray, largely

massive, but with a few scattered silty
limestone nodules 3-5 cm (1-2 in.) in
diameter, in beds 12-25 cm (5-10 in.)
thick, separated by fine-grained lime-
stone beds 2-5 cm (1-2 in.) thick, some
of which are cut into boudins by
partitions of siltstone parallel to
slaty cleavage •*««ooc.J.i 50 170

Covered interval ,,..>,.,,. . „ .. 20 70

6. Siltstone, gray, calcareous, in faintly
laminated beds separated by prominent
partings at intervals of 5-25 cm (2-10
in.); brachiopods fragmentary, scarce,
Mo nog rapt us sp. rare 20 65

122

Covered interval 25 85

5. Silt stone, calcareous, like unit 6 above 15 50

4. Siltstone and fine-grained sandstone,

calcareous, light-gray, mostly in well-
defined graded beds 12-24 cm (8-10 in.)
thick, some prominently laminated. A
dark-gray noncalcareous siltstone bed
1 m (3 ft) thick, 20 m (60 ft) above the
base of the unit contains 17 genera of
brachiopods including Cyrtia sp.,
Eospirifer cf. _E. radiatus (Sowerby) ,
Howell ella sp., Merista sp. , and
Sphaer irhynchia sp . , trilobites and
c orals 20 70

Lower Silurian Rocks, 54 m (175 ft) measured, of
which 36 m (115 ft) is exposed)

3. Pebble and granule conglomerate, sandstone,
and siltstone, calcareous, in well-
defined graded beds 12-25 cm (4-8 in.)
thick; fragments of brachiopods and
cor als 14 45

Covered interval 10 35

2. Conglomerate with rounded pebbles and

granules, mostly of quartz, but with

abundant felsitic volcanic rocks and

tabular fragments of slate and siltstone

as much as 15 cm (6 in.) long; beds poorly

defined, about 1.5 m (4 ft) thick;

scattered brachiopods including

Stricklandia lens ultima Williams ,

tabulate and rugose corals 20 65

Covered interval 8 25

1. (Beds exposed in floor of Bowlin Pond Road)
Siltstone, greenish-gray , and fine-
grained sandstone; 2 beds of granule
conglomerate, each about 25 cm (10 in.)
thick 2 5

Beds below not exposed; distance to Grand Pitch

Formation ,,,... 175 575

Return to Grand Lake Road.

17.1 Grand Lake Road. TURN LEFT.

18.2 Enter Traveler Mountain Quadrangle.

123

19.3 STOP 5. "Hurricane Deck." Overlook and exposures of

Matagamon Sandstone. For description and discussion see Stop
1 of trip B3 (Rankin, this guidebook).

Continue westward on Grand Lake Road.

20.8 Bridge over East Branch of the Penobscot River. Immediately

after crossing bridge, TURN LEFT. Road skirts campground;
keep right, away from campground.

21.0 Culverts (if luck is with us) over backwater. Woods road

follows approximately old Eagle Lake Tote Road.

26.0 STOP 6. Haskell Rock; park vehicles as space permits here and

half mile to south.

The purpose of this stop is to examine the Upper
Ordovician succession along the west side of the river between
Pond Pitch and Haskell Rock. Although fossiliferous Upper
Ordovician rocks (equivalent to the British Ashgill Series)
are known from a few other places in Maine (Lobster Mountain
Volcanics of Boucot and Heath (1969); Blind Brook Formation of
Hall (1970); and at several places in the vicinity of
Ashland) , this section has so far proved unique for its
completeness, the apparent amount of time that it represents,
the abundance and diversity of its fossils, the variety of
rocks that it includes, and its relatively simple structure.
The presence of this thick section (1,200 m, 3,500 ft) here,
compared to the total absence of equivalent rocks in the
Bowlin Brook exposures of Stop 4, three miles to the northeast
in the same outcrop belt, may be due as much to structural
causes as to stratigraphic ones. One possibility not hitherto
entertained is that faulting at this level rearranged bodies
of rock so their present distribution does not reflect their
original relations.

Figure 3 represents exposures on both sides of the river
constructed from a 1: 2,400-scale tape and compass survey made
in 1976. The exposures on the west side that should be seen
during this trip are listed below, from south to north.

1. Basalts at Pond Pitch, identified by Rankin (1961) as
olivine basalt with rare pillows, about 100 m (300 ft) thick;
quartzite of the Grand Pitch Formations crop out about 75 m
(200 ft) to the south of the southernmost basalt outcrop on
the east side of the river, but the contact between them is
not exposed. Small pillows near the top of the basalt,
immediately below the contact with the overlying sediments,
are visible when the river level is low.

2. Lower siltstone, about 25 m (80 ft) exposed; dark-gray
siltstone, bedding obscure in most places, strongly cleaved;
fossils obtained from small roadside ledge on west side of

124

CO

Q
O
0.
O

2
u

2

PQ .3

XXX

c

o

•o

o

<u

J=

cfl

a

D£

0

>>

u

rsj

*-»

m

If)

■•-»

o

a)

w

U

X X

X

1 *

X X

V)

a

a)

r

u

*-<

2

•♦->

l/l

u.

U

0

o

o

t,

Li

-n

■ — i

•*->

m

tfl

U)

ill

(h

>>

(-,

o

•♦-»

CO

CO

0)
I— «

o

o

CO

I— I
r— I

CD
■•-•

o

g

u

>>

u

CJ

CU

I— I

a

— -i

Li

h
i
i

i

i
cfl

c
cu

T3
(-,
CO

fa

X

6

»— i

X X

l

X —

ill

C ♦* > D,

o a>,o
a <u J) -o
o h u 5

£coOCQ

X X X X X x

W co

811

co a>
U tf

<p

■v.

ft N^

••

o
o
o

<r

' «• • • «

XX X XXX X

CO

3.2

cti

u
o

u

0)

1— I

o

a

CO

cfl

c
o

o

XXX

h 3 f o h

UwJWU
i

I

x

CO

i— I CU

r-i a

cu co

CO

5 •

a§

a) m

cu <*-<

!*

4-1 T3

CU CU

,fi 4J
u

d d

o Ui

•H 4-1

4-1 CO

CJ c

cu o

CO o

o «

•H >_i

00 CU

o >

O Pi

cu

O 4-1

o
o

• CO

o

cu a

M CU

3 Ph

ac

fa

o

T3

o a

a cu

O 4J

•H JJ

X -H

CO £

M o-'

PQ

CO

.Q

• O

•

V© 4J

CO

r»» «h

d

CT\ d

u

•H «0

cu

s

4-1

a

4-1

4J M-l

CO

CO O

a

3

GO ^

>•»

3 w

x>

<3 -H

CO

i

•H CU

o

^ >

x

£ -H

CO

O d

H E=>

CO

u

^!

<0 CU

o

a ^

O

CO u

u

127
TRIP B-5

PLANT FOSSILS (PSILOPHYTES) FROM THE DEVONIAN
TROUT VALLEY FORMATION OF BAXTER STATE PARK

Andrew E. Kasper, Jr.
Department of Botany, Rutgers University
Newark, New Jersey

INTRODUCTION

Although small in geographical extent, the Trout Valley
Formation contains one of the richest early- land-plant depos-
its in the world. Dorf and Rankin first described the flora
and defined the formation in 1962. Since then, Andrews (Bio-
logical Sciences, Univ. Connecticut) and his students in a
series of papers over the years have reported numerous taxa
from these strata. A recent paper with reconstructions and
illustrations of both micro- and mega-plant fossils has sum-
marized their work up to now (Andrews et al . , 1977).

The Trout Valley Formation consists of relatively unde-
formed continental strata lying in the trough of a synclinori-
um whose axis strikes east-northeast in the Traveler Mountain
Quadrangle, Maine. The formation contains the youngest sedi-
mentary rocks in the area. Their maximum exposed thickness is
about 1500 feet and they outcrop over an area 1.5 by 8 miles
in the valley of Trout Brook northwest of Traveler Mountain in
Baxter State Park (Dorf & Rankin, 1962). These clastic rocks
are a heterogeneous assemblage of shale, silts tone, sandstone
and conglomerate. The formation unconformably overlies the
Traveler Rhyolite which, in turn, rests conformably on the
Matagamon Sandstone (Dorf & Rankin, 1962) . The marine fauna
of the Matagamon Sandstone is of Becraf t-Oriskany age (=Sieg-
enian of Europe, Lower Devonian; Rankin, 1965).

Dorf and Rankin (1962) date the Trout Valley Formation
as Early Devonian based on the flora. Because of lithological
similarities, Rankin (1961) had previously correlated the for-
mation with the marine beds of the Tomhegan Formation to the
southwest. Boucot et al. (1964, p. 94) date the Tomhegan For-
mation as Schoharie (=Emsian, Lower Devonian). A palynolog-
ical analysis by Andrews et al . (1972) suggests an Emsian/
Eifelian age (late Lower/early Middle Devonian) . An absolute
date for the underlying Traveler Rhyolite is given as 360-10
m.y. (Bottino et al . , 1966) which suggests a much younger age,
Givetian/Frasnian (late Middle/early Upper Devonian) . In
their summary paper based on both plant megafossils and paleo-
palynology, Andrews et al. (19 77) suggest that the formation
is Emsian (late Lower Devonian). Finally, the recent discov-
ery of the hitherto Middle Devonian lycopod, Leclercqia
complexa, in the formation adds to the dating problem and
favors a Middle Devonian age (Kasper & Forbes, 1979).

128

The majority of species described in the Trout Valley For-
mation belong to a group of extinct plants which previously
had been called 'psilophytes'. In 1968 Banks re-classified
the psilophytes into three distinct Subdivisions, one of which,
the Trimerophytina, is abundantly represented in the Trout
Valley flora. In addition to this Subdivision, lycopods--well
known from the 'scale trees' of the Carboniferous--are also
present in the formation, however, only as small herbaceous
forms .

PSILOPHYTE CLASSIFICATION

Early land plants have been called 'psilophytes' after
the genus, Psilophyton, described by J. W. Dawson in 1859 from
the Gaspe Peninsula, Quebec. Dawson's historic Psilophyton
princeps was one of the first-reported early land plants from
North America. The name Psilophyton, i.e., naked-plant, cor-
rectly characterizes the group as plants without leaves. For
many years after, paleobotanists assigned the numerous and
varied early Devonian fossil plants to this single catagory,
the psilophytes. It was becoming readily apparent, however,
that the psilophytes were a heterogeneous assemblage and an
artificial taxon. In 1968 and more recently, 1975, Banks re-
classified the psilophytes into three major groups based on
structural and reproductive features.

Subdivision Rhyniophytina :

The Subdivision name comes from the genus Rhynia described
from the silicified Devonian peat beds near Rhynie, Scotland.
The Subdivision Rhyniophytina contains small leaf-less plants
a few decimeters in height with dichotomously forking stems
and single sporangia borne at the branch tips. The sporangia
or spore sacs are generalized as being spindle-shaped and split-
ting lengthwise. A cross-sectional view of ' rhyniophyte' stem
anatomy shows a small centrally-located strand of conducting
tissue round in outline with the first- formed cells in the
center (centrarch) .

The rhyniophytes are the most primitive vascular land
plants known- -primitive from both a botanical and geological
standpoint. Morphologically and anatomically they are the
smallest and simplest of plants. Geologically they contain
the oldest unquestionable vascular plant genus, Cooks onia .
Cooksonia was described by the paleobotanist , Lang, in 1937
from latest Silurian (=Downtonian) strata of Wales. Banks
(1974, 1975) has discussed the controversial topic of the
'oldest land plant' in two recent papers.

129

Subdivision Zos terophyllophytina :

The Subdivision Zos terophyllophytina is the second group
to which a portion of the psilophytes have been assigned. The
name comes from the early described genus, Zosterophyllum
(Penhallow, 1892) from the Lower Devonian of Scotland. The
' zosterophylls ' are larger plants than the rhyniophytes and
may have attained heights up to a meter. They also had di-
chotomously forking stems, however, 'overtopping' was common.
Overtopping occurs when one limb of a dichotomy continues as
the main shoot and the other limb is restricted in growth as a
side branch. Even though all parts of the plant are equiva-
lent botanically speaking, the plant appears to have a 'main
stem' and 'side branches ' --the pseudomonopodial habit. Over-
topping and an increase in size of the conducting tissue per-
mitted plants to grow taller. The conducting tissue in cross-
sectional view is much larger, elliptical in outline and has
the first-formed cells at its periphery (exarch) .

The key to identification of the zosterophylls is the lo-
cation of their spore sacs, sporangia. The sporangia are at-
tached along the stem--in a lateral position--rather than at
the tips or terminal position as in the rhyniophytes. The
spore sacs are kidney-shaped (reniform) and open by a distal
suture. Finally, the zosterophylls may either be unornamented
or their stems may display a variety of surface ornamentation
or enations . Enations vary from species to species and are
important in identification. They range in form from multi-
cellular spines or glands to deltoid tooth-like emergences.

Subdivision Trimerophytina :

The third group to which a portion of the psilophytes
have been assigned is the Subdivision Trimerophytina (Banks,
1968, 1975). The majority of the Trout Valley specimens be-
long to this catagory. This Subdivision contains the largest
plants of the three. The ' trimerophytes ' were probably over a
meter tall, with a robust stem in which overtopping was so
pronounced that specimens display a distinct main stem and
side branches. These plants had much more conducting tissue
than members of the previous two groups. This, along with
thick-walled outer cortical cells, supported these taller
plants. The conducting tissue was circular in cross section
and the first-matured cells were in the center (centrarch) .
The location, shape and dehiscence of the sporangia are diag-
nostic. The spindle-shaped or fusiform sporangia are located
at the ends of the side branches and form dense clusters rath-
er than being solitary as in the rhyniophytes or scattered as
in the zosterophylls. The spore sacs opened by means of a
longitudinal slit. Some of the trimerophytes were unornamen-
ted, while others bore hair-like, gland-like or spine-like
processes on their stems --important characteristics in field

130

identification

FLORA OF THE TROUT VALLEY FORMATION

Plant fossils from the Trout Valley Formation were first
described and illustrated by Dorf and Rankin in 1962. These
significant but fragmentary specimens were assigned to six
taxa of early land plants. This initial report was important
in making geologists aware of the presence of plant fossils in
northern Maine and their potential use in dating and correlat-
ing the isolated continental deposits in the region.

In 1964 and 1965 Forbes (Univ. Maine, Presque Isle),
Mencher (City College, C.U.N.Y.) and Schopf (U.S.G.S., Colum-
bus) guided Andrews to several plant localities in northern
Maine and encouraged him to study the plant fossils of the
area. In 1968 Andrews et al. reported a new species of
Psilophyton, P. forbesii, named after its discoverer.

In 1969 Gensel et al. described a new plant,
Kaulangiophyton akantha, which appears to be intermediate be-
tween the zosterophylls and the lycopods . This is an impor-
tant addition to the accumulating evidence that the lycopods
arose from the zosterophyll line.

Six years after the initial introduction to the area,
Andrews and Kasper (1970) published a summary report on the
flora. The text was accompanied with illustrations of speci-
mens and reconstructions of plants. The age of the formation,
based on plant fossils, was given as middle or upper Lower
Devonian.

In 1972 a new trimerophyte, Pertica quadrifaria, was de-
scribed by Kasper and Andrews based on exceptionally complete
compression/impression specimens. The material permitted an
accurate reconstruction of the plant providing information on
its size, arrangement of side branches on the stem and the
branching patterns in both fertile — sporangium bearing- -and
sterile branches.

In 1974 Kasper et al . described two new species of
Psilophyton, P. dapsile and P. microspinosum, besides pre-
senting additional information on the previously reported
species, P. forbesii and P. princeps . This paper illustrated
for the first time trimerophyte remains from the Fish River
Lake Formation in the Eagle Lake/ Saint Froid Lake area (Winter-
ville and Eagle Lake Quadrangles) of northern Maine.

A detailed review of the megaplant fossils along with the
first illustrated analysis of the microflora was presented by
Andrews et al. in 1977. This is the best summary article, to

131

date, containing numerous photographs and reconstructions.
Again, an age of either "...late Early Devonian or earliest
Middle Devonian...." is suggested (Andrews et al., 1977, p.
283).

Finally, Kasper and Forbes (1979) presented the first de-
tailed report of a lycopod from the formation. Leclercqia
complexa, although preserved in a fragmentary condition, is
readily identified because of the unique morphology of its
leaves. Up until now this lycopod has been known only from
the Middle Devonian, so it adds to the controversy regarding
the age of the Trout Valley Formation.

A revised list of plant mega- and microfossils is pre-
sented below. Each taxon is followed by one or two selected
references recording its presence in the Trout Valley Forma-
tion and, if illustrated, by the plate and figure numbers.
For comments on Dorf and Rankin's (1962) original determina-
tions see Andrews et al. (1977, p. 272).

Megafossils :

Subdivision Rhyniophytina

Taeniocrada sp . -- Dorf & Rankin, 1962, PI. 140, Fig.
5"; Andrews et al., 1977, PI. VI, Fig. 4.

Subdivision Zosterophyllophytina

Sawdonia ornata -- Dorf & Rankin, 1962, PI. 140,

Fig. 1-4.
Kaulangiophyton akantha -- Gensel et al . , 1969.

Subdivision Trimerophytina

Psilophyton forbesii -- Andrews et al., 1968; Kasper

et al. , 1974, Fig. 21-26.
P. dapsile -- Kasper et al., 1974, Fig. 5-9.
P. microspinosum -- Kasper et al., 1974, Fig. 13-19.
P. princeps -- Kasper et al., 1974, Fig. 28-33.
P. sp. -- Andrews et al., 1977, PI. II.
Pertica quadrifaria -- Kasper & Andrews, 1972.

Subdivision Lycophytina

Drepanophycus sp. -- Andrews 6c Kasper, 19 70, Fig. 2B;

Andrews et al., 1977, PI. VI, Fig. 3.
Leclercqia complexa -- Kasper & Forbes, 1979, PI. Fig

1-12.

132

Incertae Sedis

Prototaxites sp . -- Andrews & Kasper, 1970, p. 12
Sciadophyton sp . -- Kasper et al., 1974, p. 358.
Thursophyton sp . -- Andrews et al., 19 77, PI. VI,
Fig. 1-2.

Microfossils :

Spores (Andrews et al., 1977)

Deltoidospora sp., cf. D. priddyi -- PI. VII, Fig. 1.
Apiculiretusispora sp. -- PI. VIII, Fig. 1-2.
Emphanisporites rotatus -- PI. VII, Fig. 5; PI. VIII,

Fig. 5.
E. annulatus -- PI. VII, Fig. 6-7.
cf. Clivosisporites verrucata -- PI. VII, Fig. 3.
Tholisporites sp . , cf. T. chulus -- PI. VII, Fig. 2.
Grandispora sp., cf. G. douglas townense -- PI. VIII,

Fig. 3-4.
Grandispora sp. -- PI. VII, Fig. 4.

Chitinozoa (Andrews et al., 1977)

Sphaerochitina sp. -- PI. VIII, Fig. 6.

LOCALITY 1

Psilophyton dapsile
?Psilophyton sp .
Kaulangiophyton akantha
Taeniocrada sp .
Thursophyton sp .
Sciadophyton sp .

This locality was first reported by Dorf and Rankin in
1962 and has turned out to be the most productive fossil site
in the Trout Valley Formation. The best-preserved plant here
is the trimerophyte, Psilophyton dapsile (Kasper et al . , 1974).
P. dapsile was small, a few decimeters tall, with unornamented
dichotomous or, occasionally, pseudomonopodial (overtopped)
axes. The ultimate branchlets bore dense clusters of small
sporangia. Specimens can be identified by their 2 mm wide
axes, dichotomous branching and numerous elliptical paired
sporangia about 2 mm long.

Also abundant at this site are specimens of a much larger
plant tentatively assigned to the trimerophytes as ?Psilophyton

133

sp . (Andrews et al., 1977, PI. II). Overtopping (pseudomono-
podial habit) is quite evident in this plant with its distinct
main axis and dichotomous laterals. The stems are 3 mm wide
and are readily distinguished by the presence of longitudinal
grooves and ridges probably resulting from supporting tissues
revealed by compression. Specimens appear similar to those
illustrated by Dorf and Rankin (1962, Fig. 6, 8) as Hostimella
sp . (now: Hos tinella) and Aphyllopteris sp. Hostinella and
Aphyllopteris are both form genera- -genera which are not read-
ily assignable to any higher catagory because of the limited
information provided by the specimens. The plant in question
is referred to Psilophyton because of its overall habit. How-
ever, since sporangia have not been found as yet, a definite
assignment to this genus cannot be made.

Kaulangiophyton akantha (Gensel et al . , 1968) as recon-
structed was a small plant with horizontal and erect axes up
to 9 mm wide. Large (8 mm in diameter) ovoid sporangia on
short stalks were borne attached along the erect axes. Small
(2 mm long) spines were scattered along both prostrate and up-
right portions. K. akantha is important because of its inter-
mediate position between the zos terophylls and the lycopods .
It is tentatively classified here under the Zosterophyllo-
phytina.

Taeniocrada (Andrews et al., 1977, Fig. 4) is a com-
mon Devonian plant genus described from around the world, how-
ever little is known about the plant itself. It is easily
recognized by its broad (l%-2 cm) ribbon-like axes bearing a
central strand 2 mm wide and dichotomizing infrequently. It
is abundant at this locality.

Thursophyton (Andrews et al . , 1977, PI. VI, Fig. 1, 2) is
another genus that is not known well, botanically speaking.
In fact some authors use it as a form genus for fragmentary
remains of axes densely covered with small, delicate, spine-
like or hair-like enations . These 'leafy' axes branch pseudo-
monopodially or dichotomous ly , are about 5-6 mm wide and are
clothed with emergences about 2 mm long. They are easily
recognized but specimens are scarce. Sciadophyton has been
found at this locality only once in the many years of collect-
ing, so it too is rare.

LOCALITY 2

Pertica quadrifaria
Leclercqia complexa

This second locality was discovered by Forbes and I in
July, 1971 (Kasper & Andrews, 1972); in later publications it
is referred to as Locality # 7 (Andrews et al . , 1977). Sev-

134

eral different plants are present at this site, two of which
have been published: the trimerophyte , Pertica quadrifaria,
and the lycopod, Leclercqia complexa.

As reconstructed, P. quadrifaria was a plant a meter or
more tall with marked overtopping giving the appearance of a
distinct main stem and side branches (Kasper & Andrews, 1972).
The branches were arranged in a spiral and in four ranks or
rows 90 apart (quadriseriate) . The laterals were either fer-
tile, i.e., sporangium-bearing, or sterile. Both types of
branches were dichotomous and three dimensional. The sporan-
gia are elliptical, 2-3 mm long and were aggregated into spher-
ical masses . The specimens are easily identified by the large
main stems (l%-2 cm wide) with dichotomously forking side
branches to which are often attached the sporangial clusters.

At this locality but preserved in a very fragmentary man-
ner is the lycopod, Leclercqia complexa (Kasper & Forbes, 1979)
Identification of Devonian lycopods rests in large part on
their leaf morphology. L. complexa is distinguished from
other lycopods by its five-tipped leaves. The distal part of
the blade is divided into a long tapering median segment with
two shorter pointed segments on either side. Maceration of
rock samples in HF acid frees large quantities of nearly com-
plete leaves from the matrix. L. complexa can be recognized
at the site by examining rock specimens with a hand lens. The
highly coalified leaves are very reflective and readily show
the five distal segments. L. complexa was first described
from the late Middle Devonian Panther Mountain Formation of
New York State (Banks et al., 1972). The implications of this
as regards the age of the Trout Valley Formation are discussed
in Kasper and Forbes (19 79).

CONCLUSION

Although there is still much information to be obtained
from the plant fossils of the formation, several important
benefits have already accrued from studies to date. Botanical-
ly speaking, with the description of new genera and species
the flora has provided a better understanding of the diversity
of forms present during early Devonian times. Secondly, the
remarkable preservation of some specimens --having large por-
tions of the stem with fertile and sterile branch systems in-
tact—has given us an opportunity to accurately reconstruct
the plants in a true- to- life manner. Finally, the presence of
several species of the genus, Psilophyton, displaying a wide
variety of morphological features, has allowed us to speculate
on possible evolutionary trends within the group (Kasper et al.
1974) . It is a rather uncommon circumstance in early Devonian
paleobotany to have several species of a single genus preserved
within the limits of a geographically restricted formation.

135

Geologically speaking, accurate dating of the formation
will probably rest on the plant megafossils . The spores, un-
fortunately, are poorly preserved "... showing a high degree
of coalif ication ... as might have been expected from the high
degree of diagenesis evident from the lithology ...."
(Andrews et al . , 1977, p. 275). Other paleontological data is
scarce or yet to be discovered. Secondly, the morphological
completeness of the plants of the Trout Valley flora permit
identification and comparisons with isolated and fragmentary
remains in other deposits in the region. At present, Forbes
and I are working on two very small, geographically isolated
and fragmentary floras in Nova Scotia. We are able to make
identifications based on comparisons with the wider variety
and more completely preserved plants of the Trout Valley For-
mation. Finally, it is hoped that after a thorough descrip-
tion, the flora of the Trout Valley Formation may serve as the
basis along with other floras for a mega-plant biostratigraph-
ic scheme for the region. Such a scheme would satisfy the in-
itial request of the geologists working in the area who intro-
duced us to the plant fossils with the hope that we could date
the rocks and aid in correlating the numerous isolated inter-
montane continental deposits in the northern Appalachians.
This is our ultimate goal and, at the same time, this demon-
strates the important role to be played by the Trout Valley
flora.

REFERENCES

Andrews, H.N. and A.E. Kasper, 1970, Plant fossils of the

Trout Valley Formation: Maine Geol. Surv. Bull. 23, p

3-16.

, and E. Mencher, 1968, Psilophy-

ton forbesii, a new Devonian plant from northern Maine
Bull. Torrey Bot. Club, v. 95, p. 1-11.

, W.G. Chaloner, and K.L. Segroves , 19 72, Early

Devonian spores from Maine, U.S.A.: 24th Int. Geol.
Congr., Abstr., p. 213-214.

, A.E. Kasper, W.H. Forbes, P.G. Gensel, and W.G.

Chaloner, 19 77, Early Devonian flora of the Trout Valley
Formation of northern Maine: Rev. Palaeobot. Palynol.,
v. 23, p. 255-285.

Banks, H.P., 1968, The early history of land plants: IN

Evolution and environment, E.T. Drake (ed.), Yale Univ.
Press, New Haven, p. 73-107.

, 1974, Occurrence of Cooksonia, the oldest vascu-

lar land plant macrofossil, in the Upper Silurian of New

136

York State: Jour. Indian Bot. Soc. --Golden Jubilee Vol.
50A, p. 227-235.

, 1975, Reclassification of Psilophyta: Taxon, v.

24, p. 401-413.

, 1975, The oldest vascular land plants: a note of

caution: Rev. Palaeobot. Palynol., v. 20, p. 13-25

, P.M. Bonamo, and J.D. Grierson, 19 72, Leclercqia

complexa gen. et sp. nov. , a new lycopod from the late
Middle Devonian of eastern New York: Rev. Palaeobot.
Palynol., v. 14, p. 19-40.

Bottino, M.L., C.C. Schnetzler, and P.D. Fullagar, 1966, Rb-Sr
whole-rock age of the Traveler and Kineo Rhyolites, Maine,
and its bearing on the duration of the Early Devonian:
Geol. Soc. Amer. Special Paper 87, p. 15.

Boucot, A.J., M.T. Field, R. Fletcher, W.H. Forbes, R.S. Naylor,
and L. Pavlides, 1964, Reconnaissance bedrock geology of
the Presque Isle Quadrangle, Maine: Maine Geol. Surv.
Quad. Map Series No. 2, p. 1-123.

Dawson, J.W., 1859, On fossil plants from the Devonian rocks

of Canada: Quart. Jour. Geol. Soc. London, v. 15, p. 477-
488.

Dorf, E. and D.W. Rankin, 1962, Early Devonian plants from the
Traveler Mountain area, Maine: Jour. Paleontol., v. 36,
p. 999-1004.

Gensel, P., A. Kasper, and H.N. Andrews, 1969, Kaulangiophy ton ,
a new genus of plants from the Devonian of Maine: Bull.
Torrey Bot. Club, v. 96, p. 265-276.

Kasper, A.E. and H.N. Andrews, 19 72, Pertica, a new genus of
Devonian plants from northern Maine: Amer. Jour. Bot.,
v. 59, p. 897-911.

and W.H. Forbes, 1979, The Devonian lycopod

Leclercqia from the Trout Valley Formation of Maine
Maine Geol. --Geol. Soc. Maine Bull. No. 1, p. 49-59.

, H.N. Andrews, and W.H. Forbes, 1974, New fertile

species of Psilophyton from the Devonian of Maine: Amer.
Jour. Bot., v. 61, p. 339-359.

Lang, W.H., 1937, On the plant-remains from the Downtonian of
England and Wales: Phil. Trans. Roy. Soc. London Ser. B,
v. 227, p. 245-291.

Penhallow, D.P., 1892, Additional notes on Devonian plants

137

from Scotland: Can. Rec. Sci . , v. 5, p. 1-13.

Rankin, D.W. , 1961, Bedrock geology of the Katahdin-Traveler
area, Maine: Ph.D. Thesis, Harvard Univ., Cambridge,
Mass .

, 1965, The Matagamon Sands tone--a new Devonian

formation in north-central Maine: U.S. Geol. Surv. Bull
1194-F, p. F1-F9.

ITINERARY

Assembly point is Shin Pond Lodge, Shin Pond, on State
Route 159 about 10 miles northwest of Patten, Maine. Assembly
time is 8:30 A.M. Topographic Maps: Traveler Mountain and
Shin Pond Quadrangles .

Mileage

0.0 At Shin Pond Lodge take Grand Lake Road north out of
Shin Pond (Shin Pond Quadrangle) .

5.9 Cross bridge over the Seboeis River.

10.4 Note Forest Service Camp and side road (right) to
Hay Lake .

14.7 Cross bridge over East Branch of Penobscot River
(Traveler Mountain Quadrangle) .

15.7 Baxter State Park entrance; Horse Mountain on left.

19.0 Pass through Trout Brook Farm area.

20.0 Road turns sharply left (south); enter Trout Valley
Formation shortly.

22.9 Cross over Dry Brook.

23.4 LOCALITY 1: The Crossing—a picnic area on the right
and before the bridge over Trout Brook; park cars
here and take path along the southeast bank of Trout
Brook for about 200 yards upstream to outcrop; refer
to text for information on plant fossils.

AT NOON RETURN TO CARS and take South Branch Ponds
road to South Branch Ponds Campsite (2.2 miles south
of The Crossing) to eat lunch; AT 1:00 P.M. RETURN
TO THE CROSSING and continue itinerary; cross bridge
and continue on Grand Lake Road.

138

24.8 Cross 'first' Town Line between T6N, R9W and T5N,
R9W.

25.7 Cross 'second' Town Line between T5N, R9W and T5N,
R10W.

26.4 LOCALITY 2: Park along road, walk to Trout Brook
about 75 yards south of road; large blocks of in-
durated bluish shale with maroon- colored impression/
compression plant fossils are present in the stream
bed; the origin of this material is the south bank
of Trout Brook about 40 yards upstream.

AT 4:00 P.M. RETURN TO CARS and drive back to Shin
Pond Lodge.

END OF TRIP.

139

FIGURES

Fig. 1. Psilophyton microspinosum. The pseudomonopodial
habit, i.e., main stem and side branches, of this trimerophyte
is evident. The side branches are either fertile (sporangium-
bearing) or sterile. The sporangia are erect and borne in
clusters. Note the small spines (=microspinosum) . The scale
bar is 1 cm.

Fig. 2. Psilophyton dapsile . This small trimerophyte is
unornamented and branches dicho tomous ly . The numerous sporan-
gia are small, pendant, and in dense clusters. Plants were
about 30 cm tall; scale bar is 1 cm.

Fig. 3. Kaulangiophyton akantha . This plant as recon-
structed has horizontal and erect axes, the latter bearing
large sporangia on short stalks in lateral position--a zostero-
phyll characteristic. Because of its large axes and sporangia
it approaches the lycopods in size. Scale is 2 cm.

Fig. 4. Psilophyton princeps . This historic trimerophyte
shows pseudomonopodial branching, peg- like enations and large
fusiform pendant sporangia in clusters on side branches. Major
axes are about 1 cm wide. Reconstruction redrawn from Hueber
(1967).

Fig. 5. Pertica quadrifaria. A trimerophyte about 1
meter tall with robust main stem and three-dimensionally
branched sterile and fertile laterals. Fertile branches bear
spherical masses of sporangia. Scale is 5 cm.

Fig. 6. Leclercqia complexa leaf. Stems of this lycopod
were densely covered with small five-tipped leaves. The dis-
tal part of the blade was divided into a long median segment
and two shorter segments on either side. The circular struc-
ture is the pad of tissue to which the single sporangium was
attached--on the upper surface of the leaf. Scale is 1 mm.

140

1

141

<s*£:

6

' - **>

**»-*

142
TRIP B-6

ORDOVICIAN AND SILURIAN STRATIGRAPHY OF THE ASHLAND
SYNCLINORIUM AND ADJACENT TERRAIN

by

David C. Roy

Department of Geology and Geophysics

Boston College

The Ashland Synclinorium is a major northeast-trending tectonic feature
in northeastern Maine. As shown in Figure 1, the synclinorium lies between
the Pennington Mountain Anticlinorium to the northwest and the Aroostook-
Matapedia Anticlinorium to the east. The synclinorium terminates with a
southwest plunge in New Brunswick just northeast of the Siegas-VanBuren area
(Hamilton-Smith, 1970) and may be traced for at least 60 miles (100 km) south-
westward past Ashland, Maine. Between Ashland and the St. John River, Silur-
ian rocks predominate in the synclinorium; southwest of Ashland Devonian rocks
form the structure.

The tectono-stratigraphic significance of the synclinorium lies in its
position between a region to the west that was uplifted during the Taconian
Orogeny (essentially the Pennington Mountain Anticlinorium and terrain to the
northwest) and an eastern area which was little affected by the orogeny and
remained a deep-water (oceanic?) basin (essentially the terrain of the
Aroostook-Matapedia Anticlinorium). As described elsewhere (Roy, this guide-
book) the Silurian stratigraphy in the synclinorium is interpreted to be a
post-Taconian continental margin sedimentary prism which may be "traced"
southwestward at least to the rangely area of Maine. Here the sedimentary
rocks of the prism are at low metamorphic grade (Richter and Roy, 19 76) and
rather simply deformed. The objective of this trip is to traverse the Ashland
Synclinorium from Portage (on the Pennington Mountain Anticlinorium) to
Presque Isle and examine pre-Taconian rocks and the Silurian lithofacies of
the continental margin prism.

Stratigraphy

The stratigraphy of the region, including the synclinorium, has been
described in some detail by Roy and Mencher (1976) and will be only briefly
treated here. A more detailed map of the region covered by this trip and
Trip C-5 is provided in Figure 2. STOPS 1 and 2 are in Ordovician rocks that
were deformed during the Taconian and were subsequently the main source-rocks
for the Silurian elastics. STOPS 3-9, 11, and 12 display the principal
Silurian lithofacies. STOP 10 near Washburn shows the pre-Taconian "ribbon-
rock" that is typical of the more eastern Aroostook-Matapedia Belt. For the
purposes of describing the Silurian stratigraphy that we will examine it is
convenient to divide the region into three subareas: I, the Frenchville area;
II, the area of northward plunge of the Castle Hill Anticline; and III, the
Mapleton-Presque-Isle area. Stratigraphic sections of these areas are given
in Figures 3,4, and 5. A fourth area in the vicinity of Jemtland, Maine, to
the north, will be the focus of Trip C-5.

143

<

z
<

_i

Q.
X

z
<

z
o

>

UJ

Q

in
o

<
in

UJ

_l

<

2

<

or
o

o
o

z

4
<E
U>

I
UJ

o
UJ

<

<r uj
I- t-
(/> <

as!

z </>

UJ o

<r 7

uj .

IS

o o
z o

or <

3 i

<r

o

K
UJ

\ I

Z
<

a

z

3

z
<

@0@H iiiiDDi

>

o

Q

or
o

o
o

i-

V)

o
o

<E

O

in

>

<

o >
o

-J
o

CE

<

Z
<

3

o

2

O

z
o

<E
O

<
-I

<

<
<

2

>
or

UJ
IE

<

a.

Q <
UJ i-

<n a

3 3
I ?

s o

z
<

u

z

h-

Z

O

2

UUJ

It

a. z
z 12

3 -J

o a>

:I

o .

t» (E
.. 01 UJ
CO — I
UJ -<->

o >- *
q; ouj
3 IE 2
O
CO

o 0> .

»_J*<

W o °

a> a> a.

O Q »"

jO -

<o <

a. i i

o

o

^

o

o

*J

en

o

o

u

<t

a

u

<u

j->

w

d

QJ

.d

4-1

}-<

o

a

M-l

O

C

•-s

o

VD

•H

r^

4J

cr>

M

i— i

O

ex

a

u

d

<u

rC

"4-1

o

o

d

<u

CO

d

6

t3

d

a

d

•H

d

>>

o

o

4-1

P4

o

<D

e

4-»

o

S-i

T3

M— r

c

d

'd

0)

o

•H

•H

<4-l

toO

•H

O

•d

rH

o

O

6

<u

Nw/

60

T3

<D

N

•H

rH

03

•

U

CD

<D

d

d

•H

a)

d

o

S

0)
5-1

d

60
•H

144

\ -., y/ ; osjth

OC ROY

Figure 2: Geologic Map of the Ashland Synclinorium and adjacent
areas. (modified from Roy and Mencher, 1976)

EXPLANATION FOR FIGURE 2

Stratified Rocks

145

Z

J <
<C H
H 2
Q O
W >
S W
Q

Dm

Mapleton Sandstone
Acadian Orogeny

Dl

Sandstone, Conglomerate,
Limestone and slate

Du

Undiff. Lower
Devonian

Dd

Dockendorff Group

DSfh

Fogelin Hill Formation

DSjfh

Fogelin Hill-Jemtland Undiff.

Su

Undiff. Upper Silurian

.•::'sj'.v.

Jemtland Formation

Z
<

H

D

H
CO.

H
U
H
>
O

Q
«
O

Sff

Sfs

Sfq

Sfu

Sfc

sfg

Sns

New Sweden
Formation

Frenchville Formation
Sfg Graywacke Member
Sfc Conglomerate Member
Sff Feldspathic Sandstone Member
Sfs Sandstone-slate Member
Sfq Ouartzose Sandstone Member
Sfu Undifferentiated

Taconian Orogeny

OSar

Aroostook River
Formation

Opm

Pyle Mountain
Argillite

x *

k

Oml

Winterville
Formation

Madawaska Lake
Formation

Ss

Spragueville
Formation

OS emu

o

CO

o
3

OScml

Carys Mills
Formation

OS emu Upper Mber
OScml Lower Mber
OS cm Undiff.

Intrusive Rocks

Munson Granite

146

_Ordovician Lithofacies

Three main lithofacies characterize the Ordovician System. In the west
spilitic volcanic rocks, slate, graywacke and minor conglomerate comprise the
Winterville Formation (STOP 1). Volcanic rocks predominate in exposures of
the Winterville but sulfidic slate and graywacke are abundant in eastern
portions of the formation and may be generally more important then indicated
by outcrop frequency. Disseminated sulfides, especially pyrite, are ubiquitous
in the Winterville and massive sulfide has recently been discovered in
apparently commercial abundance near Bald Mountain west of Ashland. The
Winterville is of Caradocian-Ashgillian age but the bottom of the formation is
neither defined nor dated.

The Winterville Formation underlies much of the terrain in the vicinity
of Ashland and in the core of the Castle Hill Anticline to the east (Figures 1
and 2) . The extent of the Winterville in these areas is considerably south-
east of its general surface extent farther north, but is similar to the
eastern limit of rocks of similar age and lithofacies farther south (Pavlides,
1965, 1973). It is probable, but not as yet demonstrated, that the Ashland-
Castle Hill Winterville connects in the subsurface with the main body of
formation to the west. It is also possible that older phases (Caradocian and
older) of the Winterville lithofacies extend beneath the Carys Mills Formation
to the east and are continuous with similar rocks in the Tetagouche Group in
New Brunswick.

In the vicinity of Portage (STOP 2) the Winterville interfingers with
slate and graywacke of the Madawaska Lake Formation. The Madawaska Lake is
also of Caradocian-Ashgillian age and is considered to be largely a temporal
equivalent of the Winterville. The contact between the two formations between
Portage and Square Lake (Figure 2) is considered to represent the zone of
interfingering rather than a discrete plane of contact; it is possible that
the portion of the contact from Moose Mountain northward is a fault.

Rocks of Late Ordovician age east of the Ashland Synclinorium are found
in the Carys Mills Formation (STOPS 9 and 10; Pavlides, 1968; Roy and Mencher,
19 76). In the Caribou-Washburn area the Carys Mills has been subdivided into
two members (Roy, 1978b): the Lower Member is dominantly slate and graywacke with
lesser micrite and contains all of the Ordovician fossil localities; the
Upper Member is composed of slate and micrite in approximately equal proportions
("ribbon rock") with much lesser graywacke, and contains the Early Silurian
fossil localities. This subdivision of the Carys Mills is not recognized in
the type area by Pavlides (1976); although lithologically divisible in the
eastern part of the region of Figure 2 it remains to be shown that the members
as defined here can be carried more widely. The Carys Mills as mapped by
Hamilton-Smith (1970; this volume) in the Siegas area of New Brunswick is
considered here to be the Upper Member. In the Siegas area on the western
flank of the Ashland Synclinorium, Hamilton-Smith has shown the "ribbon-rock"
of the Carys Mills to overlie gradationally the Madawaska Lake Formation which
suggests a westward onlap of the limestone-rich lithofacies. This onlap is
concealed by younger rocks in the Ashland Synclinorium in the region of Figure 2.

Ordovician-Silurian Systemic Boundary

The transition from Ordovician to Silurian rocks in the region of Figure 2
is one of unconformity in the west and southwest and conformity in the east.

147

COLUMN
Thickness in Meters

LITHOLOGY AND FEATURES

ENVIRONMFNT

NOMENCLATURE

Green and red silty shale and slate
7g0 + with interlayered thin beds of very
calcareous, laminated fine-grained
graywacke and siltstone.

Thinly interlayered gray shale, non-
laminated silty shale, laminated
silty shale, and fine-grained cal-
careous graywacke. Medium-grained
and coarse-grained graywacke common.
Turbidite structures typical of
graywacke beds. Abundant graptolites,

760

686

211*

153

610

Medium-grained and coarse-grained
feldspathic graywacke. Red shale
with laminated ironstone present
locally .

Medium-grained to thick-bedded poly-
mictic conglomerate (75 to 80%) and
lithic graywacke (20 to 25<) with
minor pebbly and nonpebbly slate
(1 to 2%) .

\

Gray and green-gray lithic graywacke

Gray and green-gray slate thinly
interlayered with calcareous fine-
grained graywacke and siltstone.
Medium-grained and coarse-grained
graywacke common.

610

Green-gray slate; minor graywacke.

Distal
Flysch

Intermedi ate
and
Distal
Flysch

Proximal
"Coalesced
Fans"

Int ermed i at e
and
Distal
Flysch

Distal

Foge 1 i n
Hill
Format ion

Pi

3
O
U
O

H
«
.c
u
i>
0-

Jemt land
Formation

c

0

•r-i

d

E

0

u.
tl

H
^H
•f-»
•>

0
c
<v
u

Feldspathic
Sandstone
Member

Congl .
Member

Graywacke
Member

Aroostook

River
Format i on

Madawaska

Lake
Format i on

Figure 3:

St rati graphic section for the Frenchville area (I)
(from Roy and Mencher, 1976)

Figure 6 shows the distribution of the angular unconformity together with the
areas of inferred disconformable and conformable relations. Angular uncon-
formities have been observed at two sites. In Ashland at STOP 1 and angular
unconformity between sandstone and conglomerate of the Silurian Frenchville
Formation and the underlying sulfidic slate and chert of the Ordovician Winter-
ville Formation was observed during construction of a house. The unconformity
is no longer visible but the Silurian rocks are well exposed and under favorable
conditions small weathered exposures of the slate may be seen. The second place
where the unconformity has been observed is near Blackstone Siding to the north
(STOP 3 on Trip C-6) where Frenchville elastics overlie slate of the Madawaska
Lake Formation.

The zone of disconformity shown in Figure 6 is inferred from the absence
of marked angularity between Frenchville and pre-Frenchville units close to the
contact and also the absence of earliest Silurian fossil localities in the
lower Frenchville. Since the lower half of the Frenchville is not well dated
the absence of earliest Silurian localities in that formation may be more of a
sampling problem than an indication of hiatus. Therefore the area of discon-
formity in Figure 6 may best be viewed as a transitional zone between regions
where unconformity are more confidently inferred.

148

AGE

COLUMN
(Thickness in Meters)

LITHOLOGY AND FEATURES

ENVIRONMENT

NOMENCLATUR

760

^-T 1070

Thinly interlayered gray shale, non-
laminated silty shale, laminated
silty shale, and fine-grained cal-
careous graywacke. Medium-grained
and coarse-grained graywacke common.
Turbidite structures typical of
graywacke beds. Abundant graptolites,

Medium-grained to fine-grained gray-
wacke with green-gray slate and
minor micritic limestone. Graywacke
beds display turbidite features.
Pebbly graywacke and conglomerate
locally near base.

=? 1500 +

Di sconf ormi ty or Slight Angular
Unconformity

Poorly cleaved silty argillite and
green-gray slate. Silty argillite
is commonly f os s i 1 i f tous .

Andesitic, basaltic, keratophyri c ,
and rhyolitic volcanic rocks and
associated shallow intrusives;
black slate; red, green, and gray
chert; graywacke; and conglomerate.

Intermedi ate
and
Distal
Flysch

Intermedi ate

Coalesced

Fans

Submarine

Eros i on

and/or

Nondeposi tion

I

Distal
Shallow
Water

r

Submarine

Volcani sm_

and

Sedimentation

Figure 4:

Stratigraphic section in the Castle Hill area (II) .
(from Roy and Mencher, 1976)

a.

3

e

o

T) 0

kc

C -ri

o

0 *i

r-t «

B

•p a

a)

B U

JC

V o

hi

•-> &.

U

0.

c

o

m

^H

+j

4->

0

«

,-i

a

CO

u

o

T3 U

h.

C V

o) a

U

B

f-t

ti V

.H

c S

>r4

o

>

♦J

J2

m

c»

■a

C

c

n

d

h

to

b.

Taconian

Hiatus

Pyle

1 Mountain

|Argillite

Wintervill

Format ion

The large region of systemic conformity is based on gradational upward
transitions in shale-rich sections from dated Ordovician to dated earliest
Silurian units. The sytemic boundary in most places is within the Carys Mills
Formation and may be approximated by the transition from the lower to the upper
member. The projection of an area of conformity toward Ashland along the
Aroostook River is based on the presence in that area of the Aroostook River
Formation which appears transitional between the Madawaska Lake and Frenchville
formations and contains fauna of latest Ordovician or earliest Silurian affin-
ities (Figure 3) . North from Ashland a change from angular unconformity to
probable conformity occurs over a very short distance. This distance may be
shortened somewhat by the Alder Brook Fault located south of Frenchville
(Figure 2), it is taken to be one indication of possibly steep post-Taconian
submarine slopes.

Lower Silurian Lithofacies

The Lower Silurian contains three broad lithofacies: a western coarse-
clastic facies (Frenchville Formation), a medial slate facies with lesser

149

AGE

COLUMN
(Thickness in Meters)

LITHOLOGY AND, FEATURES

ENVIRONMENT

NOMENCLATURE

760

600

\

Gray, calcareous, phyllitic slate
with lesser micritic limestone and
minor graywacke, conglomerate, and
green-gray slate. Lenticular bodies
of laminated mangani f erous ironstone
with associated red and green slate.
Graywacke and conglomerate more
abundant in western part.

1150

E 3800 i

Thinly interlayered gray shale, non-
laminated silty shale, laminated
silty shale and fine-grained cal-
careous graywacke. Medium-grained
and coarse-grained graywacke common.
Abundant graptolites.

\

Thinly interbedded and inter lami nated
gray silty shale , light-gray silty
limestone, and calcareous siltstone.
with lesser micritic limestone.
Bioturbate structures are common.

y

Thinly interlayered gray slate and
argillaceous micritic limestone
with lesser graywacke. Graywacke
beds and many of the more silty
and sandy limestone beds display
turbidite features. Exposures
have a "ribbon rock" appearance.

Volcani sm
and Marine
Depos i 1 1 on

Eros ion

Intermediate
to
Distal

Flysch

Distal
Coalesced
Fans

Distal

Distal

Calcareous

Flysch

Dockendorf f

Group

Hi atus

B

•o o

C -H

«5 *J

o,

.H «

3

*-• 0

o

0 U

u

t) o

o

>-j u.

a

ja

u

a

ti

o

a.

c -rf

> II «>

*> -a «

s= t> a

> i-

co o

b.

it

.h a

r* O

■H ^-t

> *>

n or)

3 B

trO U

* 0

U b.

O.

CO

Carys

Mills

Format ion

Figure 5:

Stratigraphic section in the Mapleton-Presque Isle area (III) .
(from Roy and Mencher, 1976)

micrite and ironstone (New Sweden Formation), and an eastern calcareous bio-
turbated mudstone facies (Spragueville Formation). The distribution of these
lithofacies is shown in Figure 7. All three litho facies contain fauna of Late
Llandoverian to Wenlockian age and the facies changes between them take place
over fairly short distances. Facies changes across major folds within the
synclinorium are common (e.g. in Castle Hill Anticline) and are thus best seen
in the plunges of the folds; in some cases the facies changes take place in
directions parallel to the structural grain (e.g. in folds west of Washburn).

The Frenchville Formation is the best studied unit (Roy 19 70a, 19 73) and
is itself divided into five members. Three of the members have virtually no
slate (Graywacke Member; Conglomerate Member; STOPS 3 and 4; and Feldspathic
Sandstone Member, STOP 5) and consist of thick-bedded sandstone and conglomerate,
These slate-poor members are vertically "stacked" in the Frenchville area
(Figure 3). The Graywacked and Feldspathic Sandstone members are found only in

150

s

I aa| Angular Unconformi+4

\-^2~\ Dlsconformi+ij or slicjhl Anqular Unconformity
Conformi+u

DC ROY

Figure 6: Nature of Ordovician-Silurian systemic boundary in the region
shown in Figure 2.

151

the Frenchville area. The Conglomerate Member is much more extensive and may-
be mapped northward to the vicinity of Stockholm along the west flank of the
Stockholm Mountain Syncline (Figures 2 and 7) .

Two slate-rich members of the Frenchville lie between the slate-poor
members and the eastern and more offshore lithof acies (Figure 7) . The Sand-
stone-Slate Member (STOP 7) is the more southerly and differs from the
Quartzose Sandstone Member in containing almost exclusively lithic and
feldspathic graywacke. The Quartzose Sandstone Member is conspicuous in its
abundance of quartz arenite and graywacke; this member is pretty much confined
to the Stockholm area and represents a change in provenance rather than de-
posit ional environment.

The slate-rich Frenchville members represent the transitional facies from
the base-of-slope deposits, represented by the slate-poor members, and the
slate dominated facies farther east. The New Sweden Formation rests conform-
ably on the Upper Member of the Carys Mills and differs from the older unit
primarily in its preponderance of calcareous slate (STOP 8) and the local
presence of laminated ironstone (STOP 11). The New Sweden is interpreted to be
a generally deep-water deposit (Roy, 19 70a). In the Mapleton area the Sprague-
ville Formation (STOP 12) intervenes stratigraphically between the Carys Mills
and New Sweden formations. Eastward the Spragueville thickens at the expense
of the New Sweden Formation which it completely replaces at Presque Isle. The
Spragueville lithofacies has been mapped as far east as Fort Fairfield by
Pavlides (1968) and Roy and Mencher (unpublished data). The Spragueville con-
sists of laminated, generally very calcareous, highly bioturbated mudstone
which in the Presque Isle-Mapleton area is uncleaved. The terrigenous compo-
nent of the Spragueville is coarser than that in the New Sweden and suggests
that the source of the Spragueville material may not be from the west; alter-
nate sources for the material have not been established but eastern sources
seem likely. The bioturbation of the Spragueville facies complicates attempts
to fit it with the non-bioturbated New Sweden and Frenchville lithofacies
farther west. Bottom conditions during Spragueville deposition were apparently
suitable to support an abundant but possibly not diverse infauna, whereas
farther west bottom conditions were less supportive of sediment feeders. With
time the bioturbated/non-bioturbated facies boundary moved eastward probably
due to basin subsidence. The facies boundary does not appear to have moved
east of Presque Isle before deposition of the overlying Jemtland Formation began.

Upper Silurian Lithofacies

The three Lower Silurian lithofacies are all succeeded gradationally by
the distinctive thin-bedded flysch of the Jemtland Formation (STOP 6; Roy,
1970b). The bottom of the Jemtland Formation appears to be more or less syn-
chronous everywhere and is of Early Wenlockian age (Roy and Mencher, 1976).
The base of the formation is usually easy to determine in the field and has
provided the closest thing to a structural marker horizon in the region. Many
of the first order folds within the Ashland Synclinorium are therefore defined
by the basal contact of the Jemtland (Figure 2) . The source area for the
formation was to the west and southwest as suggested by paleocurrent measure-
ments and sandstone petrography. The petrology and sedimentology of the
Jemtland is featured in Trip C-5 of this volume.

|g Conglomerate and Lrth.t6ra4u.Qcke || i^lT^oT^V ^

iqlomerate, Lrthic Graijwacke |-__| Slate with Manqamfert
I Feldspathic Gra^uacke 1= I Ironstone Lenses

Conq

and

Lithic and Feldsporthit Sandstone

ond Slate

Calcareous Mudstone and
Siltij Limestone Dc R0Y

152

Figure 7:

Major lithofacies of the Lower Silurian. Open arrows indicate
principal source units: W, Wxnterville Formation; M, Madawaska
Lake Formation; QD, Quartz Diorite.

153

The Jemtland Formation is succeeded by the Fogelin Hill Formation
in the Stockholm Mountain Syncline (see Figure 1 of Trip C-5) . The
Fogelin Hill Formation (not seen on this trip) is a distal flysch charac-
terized by red and green slate interlayered with calcareous, laminated
siltstone and fine-grained sandstone. The base of the formation is Early
Ludlovian in age but it may range into the Early Devonian (Roy and Mencher,
1976). The Fogelin Hill is not present in the Presque Isle area; its
absence there may be due to uplift associated with the Salinic Disturbance.
The provenance of the Fogelin Hill is unclear but it is likely to have been
derived from both western and eastern sources.

Tectono-Stratigraphic History

The history of Ordovician and Silurian (pre-Fogelin Hill) sedimen-
tation in the region is summarized by the cro6S-sectional diagrams in
Figure 8. Figure 8A shows four stages in the development of the
stratigraphy along a cross-section from Portage to Mapleton. Figure 8B
shows a similar cross-section from Ashland to Perham during the Late
Silurian to illustrate the evolution of the Early Silurian clastic wedge
derived from the south (see Figure 7) that was subsequently buried by
sediments of the Jemtland Formation. The section of Figure 8B crosses
each of those of Figure 8A approximately in the Frenchville Area.

The Taconian Orogeny may be dated by the earliest detritus
deposited in the Ashland Synclinorium. Pavlides (1968) has suggested
that the conformable transition from Carys Mills to younger units
represents the first influx of sediment from Taconian uplands. This
transition is approximately graptolite Zone 19 of the Llandovery. The
turbidites of the Aroostook River Formation are inferred to also
represent this initial detritus (Frame 2 of Figure 8A) . The Taconian
Orogeny produced uplift, mild folding, and sub greens chist grade alter-
ation of the Winterville terrain in the western and southwestern parts
of the area of Figure 2 (Richter and Roy, 1976). The Taconian effects
in this region were late aspects of a longer period of deformation
that closed an ocean basin between Portage and Quebec City (St. Julien
and Hubert, 19 75) and moved the continental margin of North American
to the area of Figure 2 (Roy, 1976; this volume). Therefore the
successive frames of Figure 8A illustrate the evolution of a Silurian
continental margin clastic wedge.

The eastern margin of the Taconian upland was irregular during the
Llandovery with a large embayment between Ashland and Portage (Figure
7). The distribution of Frenchville subfacies in the Frenchville area
suggest a substantial coarse-debris source to the south. This source
produced quartz diorite detritus (Feldspathic Sandstone Member) for a
period of time. By Jemtland time the irregularity of the Taconian up-
land margin was somewhat reduced due to Late Silurian transgression
onto the more subdued and submerging land area (Roy, this volume).

154

o

o

<f> r-

< Z

z
o

t-

UJ

_l

Q.

<

z
<

X
LU

en

<
o

Sxl

5

I
o

Z
UJ

or

z

x
*:
o
o

h-

K

LU
O
<

uj O
* 0.

evj

ro

CD

rC

4-1

M-l

o

>>

x:

•

CX. rH

cd

QJ

5-i

>

00

<U

•H

.H

4-1

cd

cd

s-4

CD

4J

CO

CO

CO

c

4-1

cd

C

•H

CL>

^

co

a

01

iH

u

•H

Cu

C/3

CD

T3

e

C

cd

o

•H

C

4J

cd

CJ

■H

<U

CJ

CO

•H

>

x:

O

o

-d

cd

>-i

QJ

O

y-i

QJ

o

X!

4J

a)

M-l

•H

O

rH

a

i-H

o

cd

•H

4-»

4-1

c

D

O

•

tH

N

CN

O

•H

>

S-i

0)

0)

o

u

x:

D

a)

00

,c!

5-i

•H

4-1

1*

00

Oh

c

c

:=>

•H

•H

4->

CO

cd

•

cd

M

CN

4J

T3

CO

CL>

<D

0

>-J.

4-J

T-\

D

cd

H

00

C

•H

•H

00

fe

•H

CO

CO

1=1

C

0)

o

•H

T3

•H

4J

C

CO

a

S

C

QJ

o

o

CO

X!

•H

1

CO

4-1

CO

cd

CO

cd

6

o

a>

S-i

u

S-i

o

U fl h

00

U

00
•H

155

References

Berry, W.B.N. , 1960, Early Ludlow graptolites from the Ashland area, Maine:
Jour. Paleontology, v. 34, pp. 1158-1163.

Boucot, A.J.; Field, M.T.; Fletcher, R. ; Forbes, W.H.; Naylor, R.S.; and

Pavlides,L., 1964, Reconnaissance bedrock geology of the Presque Isle
quadrangle, Maine: Maine Geol. Survey Quad. Mapping Ser., no. 2, 123p.

Hamilton- Smith, Terence, 1970, Stratigraphy and structure of Ordovician and
Silurian rocks of the Siegas area, New Brunswick: New Brunswick Dept.
Nat. Resources Mineral Resources Br. Rept. Inv. 12, 55p.

Miller, R.L., 1947, Manganese deposits of Aroostook County, Maine: Maine
Geological Survey Bull. 4, 77p.

Neuman, R.B., 1968, Paleogeographic implications of Ordovician shelly fossils
in the Magog Belt of the northern Appalachian region, in Zen, E-an;
White, W.S.; Hadley, J.B.; and Thompson, J.B., eds. , Studies of Appala-
chian geology, northern and maritime: New York, Interscience Pubs., Inc.,
chap. 3, pp. 35-48.

Pavlides, Louis, 1965, Meduxnekeag Group and Spragueville Formation of

Aroostook County, northeast Maine: U.S. Geol. Survey Bull. 1244-A,
pp. A52-A57. ,

1968, Stratigraphic and facies relationships of the Carys Mills Formation,

northeast Maine and adjoining New Brunswick: U.S. Geol. Survey Bull. 1264,
44p.

1973, Geologic map of the Howe Brook quadrangle, Aroostook County, Maine:

U.S. Geol. Survey Geol. Quad., Map GQ-1094.

1976, Discussion, Ordovician and Silurian Stratigraphy of Northeastern

Aroostook County, Maine; in Page L.R. , ed., Contributions to the Strati-
graphy of New England; Geol. Soc. Amer. Mem. 148, pp. 53-55.

Richter, D.A. and Roy, D.C., 1976, Prehnite-pumpellyite facies metamorphism
in central Aroostook County, Maine: in Lyons, P.C. and Brownlow, A.H.,
eds., Studies in New England geology, Geol. Soc. America Mem. 146,
pp. 239-261.

Roy, D.C., 1970a, The Silurian of northeastern Aroostook County, Maine (Ph.D.
thesis): Cambridge, Massachusetts Institute of Technology, 483p.

, 1970b, Variations in a thin-bedded turbidite sequence from the Silurian

of northeastern Maine: Geol. Soc. Amer. Abs. with Programs (Ann. Mtg.),
v.2, no. 7, pp. 669-670.

1973, The provenance and tectonic setting of the Frenchville Formation,

northeastern Maine: Geol. Soc. America Abs. with Programs, v. 5, p. 214.

1978a, Tectonic history of northeastern Maine; Geol. Soc. Amer. Abs.

with Programs (Northeast Section Mtg.), v. 10, no. 2, p. 83.

156

1978b, Geologic map on a portion. of northeastern Aroostook County Maine;

Open File Map (1:62,000), Maine Geol. Survey.

and Mencher, E. , 1976, Ordovician and Silurian stratigraphy of north-
eastern Aroostook County, Maine, in Page, L.R., ed. , Contributions to
the stratigraphy of New England: Geol. Soc. America Mem. 148, pp. 25-52.

St. Julien, Pierre and Eubert, Claude, 19 75, Evolution of the Taconian orogen
in the Quebec Appalachians: Amer. Jour. Sci., V.275-A, pp. 337-362.

Ziegler, A.M., 1965, Silurian marine communities and their environmental
significance; Nature, v. 207, pp. 270-272.

Itinerary

Mileage

0 Assembly point for the trip is in the parking lot adjacent and south

of Judd's Store at the intersection of Routes 163 and 11 in Ashland.

Starting time is 8:00 A.M. Drive north along Route 11 through "down-
town" Ashland.

.3 Intersection with Route 227 ("State Road") at the light in Ashland.
Turn left (west) and continue along Route 11 toward Portage.

1.0 Bridge across the Aroostook River.

1.1 Intersection. Turn right and follow Route 11 north.

2.5 Bridge over Little Machias River.

2.6 Intersection with Wrightville Road (unpaved) ; continue north on
Route 11.

3.15 B & A Railroad crossing.

6.9 Entrance to Pinkham's Mill. This mill is one of the largest saw mills
east of the Mississippi and it is well worth the time to take the
conducted tour.

11.0 Main intersection in Portage. Continue north on Route 11.

11.8 STOP 1. Park in turn-out at the edge of the golf course. Exposure is
part of an old road metal pit that has been somewhat "groomed".

157

Black sulfidic slate and volcanic rocks of the Winterville Formation
are exposed here in the Portage Anticline of the Pennington Mountain
Anticlinorium. Cleavage in the slate is parallel to bedding and grap-
tolites are common. W.B.N. Berry assigns the graptolites to Zone 12
(Climacograptus bicornis Zone) of the Ordovician. It is unclear at
this locality whether the basaltic rocks are intrusive or extrusive.
The volcanic rocks here and elsewhere in the Winterville Formation
have not been studied in detail. These rocks are inferred to have
been deformed during both the Taconian and Acadian orogenies. They
are presently at Prehnite-Pumpellyite grade. On a clear day there is
a nice view southeastward across the Ashland Synclinorium (low terrain)
toward the Castle Hill Anticline with Haystack Mountain on the south
end of the axial ridge which is cored by Winterville Volcanic rocks.

Turn vehicles around and head south on Route 11.

12.6 Main intersection in Portage. Continue south on Route 11.

12.85 Turn left (east) onto small gravel road. Access road appears to be
part of a driveway for a small green house but continues beyond the
house, along the edge of a field, and into the woods.

STOP 2. Park in the road metal pit at the end of the access road.

This pit, and another to the north, expose typical slate of the
Madawaska Lake Formation. These rocks are interpreted to inter finger
southwestward with volcanic rocks of the Winterville Formation. The
slate is typically gray to olive-green in color, fracture cleaved,
and very fine-grained. Orange weathering, laminated and cross-
laminated fine-grained sandstone or siltstone are usually present as
thin beds. More massive dark gray lithic graywacke is locally
abundant. Here layers of volcanoclastic sandstone and lithic tuff
are included in the Madawaska Lake. Along the west wall of the pit
is an apparently graded lithic tuff bed with pumice/scoria fragments
near its top and in the overlying slate.

Return to Route 11 and head south to Ashland.

23.8 Bridge across the Aroostook River.

24.5 Intersection of Routes 11 and 227 in Ashland at stop light; continue
straight on Route 227 heading east.

25.0 Left (north) on Cottage Hill Road near water tank.

25.2 STOP 3. Park on the right side of the road leaving spaces for the

driveways. The outcrops for this stop are on the property of Mr. and
Mrs. Judson Holmquist.

During the construction of the Holmquist house the Taconian uncon-
formity was exposed in the east wall of the basement excavation as
depicted in Figure 9. The unconformity (now covered) is between
sulfidic slates and lesser chert of the Ordovician Winterville
Formation and the overlying Conglomerate Member of the Silurian

158

\$

T~T

</!—•» —

o

z

-J

X
CO

13

d cu

CO ,C

4J

4-J

H CN

•H O

3 CU

£1 TJ J-i

w 0) 3

4-> bO

<U cO -H

(0 O [n

3 -H

CO c

x: c -h

•H

4-1 CO

co <u cd

•H M

3 (D T3

cr a)

6 co u

H QJ !fl

O -H C

SUM

•H -H

OH 111

,d co cu

4-1 U 13

O

^H 01

o c

4-4 i-H O

•H «H

£ CO 4J

O CO CO

•hob

4J fn P

CO O

> fe

CO •

O /^N

X o •

^ § "*

4-1 -H >>

C ►"} -M

CU iH

e c ih

<U O CO

CO 4-1 O

CO iH O

£> U r-{

CO

(HUH

& -H

4-) • CO

co co

•OHO

C S «w

CO

T3 4-1

■O c n)

a co

CO CO

H • -H

>

JH Jd

CO S en

<

o

>, PM

J

•4-1,0 0

O H

T3 CO

CO CU

D- C •"

CO £ ex

B o co

B

o >^

•H rH X)

&0 !-i C

o a) co

iH B -H

O >-i ,G

<U O CO

O M-4 <tj

cr.

<u
u

00
•H

159

Frenchville Formation. Both shelly and graptolite fauna were recovered
from the Ordovician slate. The graptolites suggest to W.B.N. Berry
(Roy, 1970a) an age within the span of zones 12 to 15 of the
Ordovician; the shelly fossils have been provisionally assigned to the
Ashgillian of the Ordovician by Neuman (1968, p44) . No dateable fauna
have been extracted from the Frenchville here but a shelly collection
from Locality 30 (Figure 9) is of Late Llandoverian-Wenlockian age
according to A.J. Boucot (Roy, 1970a). The Frenchville conglomerate
here is polymictic with provenance dominated by the Winterville Form-
ation. Lenses of lithic graywacke and arenite are common. Limestone
clasts and shell debris are more common here than elsewhere in the
conglomerate member and form part of the evidence for the conclusion
that the member is here of more shallow-water origin than farther north,
A major north-south fault is present just west of the house, at or near
the base of the escarpment (Figure 9). In the lowlands to the west of
the fault, one finds only Late Silurian and Early Devonian units
represented.

We will make a U-turn (carefully) and return to Route 227.

25.4 Turn left onto Route 227 and continue east.

25.85 Poor exposures of Lower Devonian sandstone at curve in road. The

Devonian rests on Silurian and Ordovician rocks unconformably in the
Ashland area. The unconformity represents the Salinic hiatus. The
lower most Devonian contains abundant limestone, lithic sandstone,
polymictic conglomerate, and monomicitic mudstone .conglomerate
(with clasts of the Jemtland Formation or volcanic rocks).

27.65 Roadcut exposure on the right (east) of basalt of the Winterville
Formation.

28.60 STOP 4. Park on road shoulder and in a small "turn-out" on the east
side of road. Walk east across small field to woods line; outcrop is
a few tens of feet in woods. We are here on the property of Mrs.
Mabel Berry.

This is the best easy-access exposure of conglomerate of the French-
ville Formation. These beds are a few hundred feet stratigraphically
above the presumed unconformity between the Frenchville and Winter-
ville formations. These massive roundstone pebble and cobble
conglomerates are thought to represent material transported down a
submarine slope into relatively deep water. The conglomerates show
little or no evidence of current sorting and commonly show a more or
less "open" texture suggestive of rapid sedimentation. Lesser lithic
arenite is also present here.

Returns to cars and continue north on Route 227.

28.75 Road ditch exposures on the right (east) near and down hill from the
power pole consist of pebbly mudstone of the Frenchville that is
stratigraphically above the beds at STOP 4. These exposures have
yielded Late Llandoverian brachiopods.

160

29.05 Bridge across Alder Brook. We here cross the Alder Brook Fault which
is a major east-west fault that extends eastward well into the
Presque Isle Quadrangle.

29.90 STOP 5. Park in turn-out just west of large road cut on right (south)
side of highway. Walk to exposure.

This is the "type" exposure of the Frenchville Formation as established
by Boucot and others (1964). The exposure is located below the site
of the former Frenchville church and lies on the west flank of the
broad Frenchville Syncline (Figure 10) . These sandstones are typical
of the Feldspathic Sandstone Member of the Frenchville. Notice the
near absence of current-produced structures, homogeneous distribution
of lithic and fossil fragments, massive bedding, and apparent graded
bedding in some of the massive beds. Figure 11 gives a generalized
stratigraphic section for this outcrop. The sandstones contain
brachiopods from three of Ziegler's (1965) depth communites (Eocoelia,
Pentamerus, and Stricklandia) which are taken to indicate downslope
transport of both sand and shell material to at least the depth of the
deepest cummunity (Stricklandia; 30-60 meters) . The Winterville does
not appear to be the major source for the detritus here; it is inferred
that a Taconian quartz-diorite like the Rockabema Quartz Diorite near
Shin Pond to the south was the source of the material.

Return to cars and continue north on Route 227.

32.7 Castle Hill Townline.

34.6 Small store on left (north) side of road. Usually soda and munchies
are available.

34.9 Demarchant Brook.

35.0 STOP 6. Park in grassy field at top of the hill on the right (south)
side of road. This may also serve nicely as a LUNCH STOP.

This is the best outcrop of the Jemtland Formation that we will be
near today. The sedimentology of this formation is the focus of
Trip C-5 where larger exposures in its type area will be examined.
This exposure is typical of the formation both in its lithology and
in the abundance of graptolites which Berry (1960) has fully described.
These beds are assigned to zone 33 of the Silurian (Early Ludlovian) .
Note the three types of shale, preservation of delicate silt laminae,
ripple cross-laminae and locally preserved ripples, micro load-cast
features associated with the siltstone laminae, and convolute bedding.
Rusty orange weathering (typical of the formation) is do to oxidation
of iron in pyrite and ankerite (?). This formation overlies the
Frenchville Formation gradationally and is interpreted to represent
deposition in deep, relatively quiet water.

Return to cars and continue east on Route 227.

35.25 From the crest of the hill here we have a nice view eastward of a line
of hills that marks the core of the Castle Hill Anticline. From south

161

hf pre * 2000'

\DCR0Y|

Figure 10:

Geologic map of Frenchville area showing the location of
STOP 5. Dotted lines outline cleared fields; base map
uncorrected airphoto. Formations are designated as in
Figure 2.

162

10-

ISO-

W>-

110-

50-

150-

->---»^4«v *~-~. **>■*>

*o-

30-

20-

110-

110-

l»o-

o-J

^J

<*>-

So-

il'' ......

II

III

IV

Covered

Figure 11: Stratigraphic

column for section
exposed at STOP 5. Rock types:
I, light-gray, medium-grained
feldspathic graywacke locally
containing concentrations of
pelite fragments; II, light-gray,
fine-grained laminated and non-
laminated feldspathic graywacke;
III, light-gray, coarse-grained
lithic graywacke with abundant
pelite fragments; IV, shale beds
and/or large pelite fragments.

to north the hills are Haystack Mountain (conical in shape) , Pyle
Mountain, McDonald Mountain and Castle Hill. Our next stop is in
the area of northward plunge of the Castle Hill Anticline.

163
36.7 Turn left (north) just beyond small church.

36.85 Turn left (north) onto gravel road. This road takes us more or less
along the crest of the Castle Hill Anticline.

38.35 Steep descent here is down-dip in the sandstone-slate member of the
Frenchville Formation.

38.95 Wade Townline.

40.3 STOP 7 . Park on grassy turn-out on left (north) side of road at sharp
bend and walk through field northward to the Aroostook River. Care-
fully descend to the river bank and walk upstream about 1000 feet to
the large exposure (see Figure 12).

The sequence exposed here (Figure 13) is transitional between the
slate-poor sandstone/conglomerate western phase of the Frenchville
and the slate dominated New Sweden Formation to the east. This
sandy-flysch phase of the Frenchville has been separated out as the
Sandstone/slate Member and is interpreted to represent an intermediate-
to-distal submarine fan sequence. The graywacke beds show many of the
classic features of turbidites and with care more or less complete
turbidites may be sampled. This is fossil locality 43 of Roy and
Mencher (1976) and it has produced conodonts which Gil Klapper places
in the interval of C~-C, of the Late Llandoverian and abundant
Art riot ret a brachiopods studied early by R.B. Neuman and more recently
by W.H. Forbes who has been extensively "mining" the limestone beds
for fossils since he and I made the first collections in the late
1960's.

Return to the cars and continue eastward.

41.3 STOP 8. Park on shoulder of road. Outcrop on right side of road.

The low exposures show typical phyllitic slate of the New Sweden
Formation which is the eastern equivalent of the Frenchville we have
been looking at. Micritic limestone beds, locally lenticular due to
bedding transposition by cleavage, are generally present in the
formation.

Continue eastward.

42.5 STOP 9. Park carefully on the shoulder of the road.

This large exposure shows the slate and graywacke typical of the
Lower Member of the Cary's Mills as 1 have mapped it in the Caribou
Quadrangle. This section faces eastward and gives way vertically to
the "ribbon rock" Upper Member of Cary's Mills seen in exposures on
the river bank just east of the outcrop here. Graptolites of Late
Ordovician (Ashgillian) age were reported from here prior to my work
and confirmed by W.B.N. Berry at my request. I have not been able to

duplicate the collection, however, and would appreciate any finds you
can make. Graptolites of Zone 18 of the Llandoverian have been found
at the "ribbon rock" exposure on the river just down-stream from here,
thus indicating that these beds are Zone 18 or older.

164

Figure 12:

Geologic map of the Castle Hill area showing the location of
STOP 7. Dotted lines outline cleared areas; base is on
uncorrected airphoto. Formations are designated as in
Figure 2.

165

?x? iz

CM

wt

H6-

20-

51

»

Medium- and Fine- a rained Lrfhic Grauwacke

m

^

nonstruc+ured qraded laminated Crosv %tmple

laminated ripples

5lQfe Lim.stone

*
m

•a

_C_

HI S9

90

S^E

) z

a

2*

nea

S)

o

1

'■;i;i;i','

'i'i'i'i'

O Fossils

Figure 13: Stratigraphic

section at STOP 7
showing flysch character of the
Sandstone/Slate Member of the
Frenchviile Formation. Fine-
grained limestone beds and
lenses are typical of those
found in the contemporaneous
New Sweden Formation nearby to
the east. Interval 23 shows
disturbed bedding caused by
slumping together with upward
diapir-like "intrusion" by sand
from bed 22.

t±=r

Si9

1G6
Return to cars and continue eastward.

43.15 Intersection. Continue straight (to the east).

45.70 Bridge over the Aroostook River at Bugbee just south of Washburn.

45.80 Turn left just east of bridge and park after making a U-turn. Descend
to the river and exposures along the east bank.

STOP 10. These exposures display typical "ribbon rock" limestone and
slate of the Upper Member of the Cary's Mills. All of the Silurian
fossil localities in the Cary's Mills of this region come from the
Upper Member of the formation and from rocks such as these. Micritic
Limestone is generally at least 50% of the unit; the slate is
typically medium- to-dark gray and quite calcareous. The limestones
are generally impure with at least 20% or so argillaceous minerals
and quartz. The impurities may be graded but are more commonly con-
centrated along 'mica-parting" laminations and cross-laminations
which show that currents were important in the deposition of the
limestone beds. It is likely that most of these limestone beds are
deep-water calcisiltites deposited by turbidity currents as suggested
by Pavlides (1965, 1968); some may be "pelagic" oozes. The more rust-
weathering beds are probably ankeritic as suggested by Pavlides and pos-
sibly more pyritiferous. Detailed studies of this "ribbon rock" have
not as yet been done, but are important (and made difficult) when the
lateral extent of this lithofacies is considered.

Return to cars.

45.9 Intersection at east end of bridge; turn right and head west.

46.5 Bear left (south) at intersection.

49.45 Intersection with Route 227. Turn right (west).

50.30 Exposure on left (south) of basal Jemtland Formation preserved in a
tight syncline and underlain by the New Sweden Formation. View west
of Castle Hill.

52.20 Turn left (south) on Turner Road. Turner road parallels the axis of
Castle Hill Anticline along its east flank.

53.20 Good view of Pyle Mountain to the right front.

54.15 Small pit on the left (east) side of road contains red slate of the
lower most New Sweden Formation. A thin calcarenite lens at this
locality produced a Late Llandoverian fauna. The lens has now been
more-or-less completely removed.

54.60 Intersection of Turner and Dudley roads. At this intersection stood
the Pyle School near which Ashgillian trilobites have been recovered
from an argillite which is now named after Pyle Mountain. Turn left
(east) on Dudley Road.

55.65 Turn right onto farm road and drive into a small road-metal pit. We
are here the guests of Mr. and Mrs. Delance Lovely.

167

STOP 11. This relatively fresh pit exposes red and green slate of the
New Sweden Formation and a basal mudstone conglomerate of the Jemtland
Formation. An approximately six-foot interval of laminated, mangani-
ferous ironstone, is exposed within the New Sweden. Such ironstone
deposits are widespread both areally and stratigraphically within
the formation and collectively form the so-called Aroostook manganese
deposits. The Dudley deposit, described as one of the richest in
manganese by Miller (1947), is along strike from this pit to the
north across Dudley Road. The red and lesser green slates are common
in the vicinity of the thicker laminated ironstones and are undoubtedly
genetically related. The varigated slate replaces the calcareous
phyllitic slates (STOP 8) that characterize most of the formation.
The basal mudstone conglomerate contains fossiliferous limestone
cobbles and pebbles; near here the conglomerate has yielded Late
Silurian (Ludlovian) brachiopods.

Return to Dudley Road.

55.85 Intersection of farm road and Dudley Road; turn right and head east-
ward.

56.80 Intersection of Dudley Road and Route 163. Continue eastward on
Route 163 toward Mapleton.

57.40 Railroad overpass in Mapleton.

58.50 STOP 12. Park on shoulder of highway; exposures on the right (south)
side of highway.

This is a typical exposure of the Spragueville Formation which is the
eastern-most temporal equilvalent of the New Sweden and Frenchville
formations. The Spragueville consists of laminated highly calcareous
mudstone. Light-gray laminae are typically 50% or more carbonate with
quartz and muscovite/chlorite being the principle insoluable material.
Microfossils are common in these calcareous laminae. This formation
is extensively bioturbated with abundant interstratal and intrastratal
burrows. Bioturbation locally completely disrupts lamination but
generally the original laminated character of the deposit is clearly
evident. Here the Spragueville underlies the New Sweden Formation
which thins to extinction near Presque Isle where the Spragueville
lies conformably between the Cary's Mills and Jemtland formations.
The Spragueville lithofacies extends into Canada east of Fort Fair-
field but has not everywhere been separated from the Cary's Mills.
Also of interest here is a poorly exposed dike of teschenite, an
analcime-rich diabase.

This is the last stop. Continue east to Presque Isle on Route 163.

59.70 Cuts of Early Devonian volcanic rocks; part of the Hedgehog Formation
of the Dockendorf Group.

61.80 Road cut on the left (north) side of the highway of the Middle

Devonian Mapleton Formation. These conglomerates are part of the

168

oldest clearly post-Acadian molasses in the northern Appalachians.
The Mapleton rests with a nearly 90 degree angular unconformity on
rocks of the Acadian-deformed Dockendorf Group and contains clasts of
most of the pre-Acadian units in region.

64.2 Intersection with Pleasant Street in Presque Isle. Turn right.

64.6 Intersection of Route 163 and U.S. 1 in downtown Presque Isle. Turn
right (south) and find your own way from here on.

169

TRIPS B-7 and C-2

BEDROCK GEOLOGY OF THE PRESQUE ISLE AREA

Richard S. Naylor*

Northeastern University, Boston, Massachusetts

INTRODUCTION

This trip includes stops in the Presque Isle, Caribou, and Ashaland
15-minute quadrangles. The author mapped the north-central part of the Presque
Isle quadrangle in 1960 as an undergraduate thesis at the Massachusetts Institute
of Technology and returned in the summers of 1961 and 1962 to extend reconnaissance
bedrock mapping to the west and north.* Although maintaining an active interest
in the area and visiting it annually, the author must stress that nearly twenty
years have passed since he last mapped in this area.

The bedrock geology of the Presque Isle Quadrangle was published by
Boucot, Field, Fletcher, Forbes, Naylor, and Pavlides (1964). That paper out-
lines the previous history of geologic mapping locally. Briefly, the area has
long been known for its rich and abundant fossil localities and many notable
paleontologists have published articles about the area. Bedrock mapping lagged
however. The pioneering work of H.E. Gregory (Williams and Gregory, 1960) produced
a useful lithologic map, but the work was not stratigraphically or structurally
oriented. Geologists from the Maine and United States Geological Surveys mapped
parts of the area in the 1940' s but were concerned chiefly with the so-called
Aroostook County manganese deposits and their immediate geologic setting. The
project on which the author and others worked in the early 1960's under the
supervision of A.J. Boucot was thus the first general purpose stratigraphic and
structural mapping of this interesting area.

D.C. Roy, Ely Mencher, and co-workers began their extensive mapping
project immediately after the end of the phase of work described above. Their
results, mostly in areas to the north and west of the Presque Isle quadrangle
are summarized in the introduction to this Guidebook and in the articles by
Roy and are not repeated here. Their stratigraphic nomenclature is followed in
this article.

GENERAL GEOLOGY

The Presque Isle quadrangle exposes four major litho-tectonic rock units
that extend to other parts of the Appalachians. In no other area are all four
units so closely juxtaposed. Interest is further heightened by the richly
fossilif erous character of most of the units in this area, providing age assignments
that can be usefully extrapolated to other less-f ossilif erous areas. A fifth
unit, the Mapleton Formation, appears to be unique to the Presque Isle area, and
is indispensible in pinpointing the age of the early phase of the Acadian Orogeny.
The five groups are outlined below, and are shown in Figure 1 (adapted from Roy
and Mencher, 1976; Boucot, and others, 1964).

*Work supported by the Maine Geological Survey.

Geology of the Presque Isle area showing lithotectonic
units described in text, with field trip route and stops.

170

UNIT 1
UN/T 2
UNIT 3

o
a

o

a

r-

/>tt l e s

171

Unit 1 comprises volcanic and sedimentary rocks exposed in the Castle
Hill Anticline in the northwest corner of the Presque Isle quadrangle and in the
Ashland quadrangle to the west (stops 1 and 2). The volcanic rocks are a bimodal
suite comprised mainly of greenstone (spilite) and sodium-rich felsite (Keratophyre)
These are closely associated with sulfidic, graptolite-bearing shale, red or black
chert, and green argillite. These rocks are of Middle and Upper Ordovician age.
This belt of rocks extends (with some gaps) for hundreds of kilometers to the
southwest and probably links with the Ordovician metavolcanic and metasedimentary
rocks of the Bronson Hill Anticlinorium. Exposures die out only a few kilometers
to the north, but the belt is postulated to continue some distance further north
below the surface (Hamilton-Smith, 1970) . The author and others have speculated
that this belt is one of perhaps half a dozen volcanic arcs significant in the
evolution of the Northern Appalachians.

Unit 2 is the Carys Mills Formation containing interlayered argillaceous
micrite, calcareous slate, and calcareous quartzo-f eldspathic gray wacke. Locally
known to geologists as "the ribbon-rock," Unit 2 is part of the Aroostook-
Matapedia litho-tectonic unit. The unit extends northeastward from Presque Isle
to the tip of the Gaspe Peninsula, but to the southwest, exposures die out in
about 50 kilometers. These sparingly folliliferous rocks are mostly Early
Silurian (Llandoverian) but are locally as old as Upper Ordovician (Ashgillian)
and possibly older.

Rast and Stringer (1974) have suggested that Unit 2 was deposited on
the floor of a small ocean basin whose eastern margin was subducted under the
volcanic are of the Tetagouche belt in New Brunswick. At the time of the authors
work, the relationship between Units 1 and 2 was unkown. Although closely juxta-
posed in the Presque Isle area they are everywhere separated by exposures of
Unit 3. Field relationships to the north and west indicate that Unit 2 was
deposited conformably on the Madawaska Lake Formation, which in turn was deposited
on the rocks of Unit 1 (Roy and Mencher, 1976).

Unit 3 is a complex of Silurian clastic formations. These range in
age from late Early Silurian (Late Llandoverian) to Upper Silurian (Wenlockian,
Ludlovian and locally Pridolian) . D.C. Roy (see Introduction) has worked out
the complex facies relationships within this Unit. He interprets it as
representing a series of submarine fans derived substantially from a source-
area developed on Unit 1 and deposited in a basin opening eastward. The contact
between Unit 1 and Unit 3 is an angular unconformity with a hiatus spanning
much of the Early Silurian. The contact between Unit 2 and Unit 3 is gradationally
conformable with no discernable structural break. The stratigraphic relationships
demonstrate the Units 1 and 2 were juxtaposed in their present position no later
than the Late Landoverian. Rocks more or less comparable to those of Unit 3 are
exposed for hundreds of Kilometers southwest of Presque Isle as the formations
comprising the western flank of the Merrimack Synclinoriu.

Unit 4 comprises volcanic and sedimentary rocks of Early Devonian
(Upper Gedinnian) age. The volcanic rocks (Hedgehog and Edmunds Hill Formations)
are distinctly more felsic than those of Unit 1. The Edmunds Hill is one of the
few andesite units described in New England, although this assignment should be
verified by modern chemical analyses. The volcanic rocks of the Hedgehog Formation
are described as andesite, trachyte, and rhyolite (Boucot, and others, 1964). The
volcanic rocks are overlain by and interfinger with the Chapman Sandstone which
is also Upper Gedinnian.

172

The interf lingering is especially well-displayed at Edmunds Hill
(Stop 8) , where a thin tongue of Chapman Sandstone is underlain by Hedgehog
Volcanics and overlain by Edmunds Hill Andesite. Fossils in the Chapman are a
mixture of plants and coarse-ribbed shells suggestive of a high-energy, near-
shore beach environment. At the Grindstone locality 9km south of Edmunds Hill,
the Chapman contains more delicate, fine-ribbed shells of the same age as the
fauna at Edmunds Hill, but suggestive of a low-energy depositional environment.
There too the water also must have been shallow because a rich assemblage of
plant fossils is also found. Further south, The Chapman Sandstone grades into
the fine-grained Swanback Slate which may be a deeper water deposit. The
relationships suggest a volcanic island in the Mapleton area, with a beach at
Edmunds Hill, and a basin deepening to the south. (A visit to the Grindstone
locality is highly recommended as part of a tour of the Presque Isle area, but
would be too time consuming to include in the present trip. The locality is
named on the topographic sheet and may be approached from the East Chapman
Road) .

Units 3 and 4 are separated by a disconformity (attributed to the
Salinic Disturbance; Boucot, and others, 1964) with a hiatus spanning the Latest
Silurian and Earliest Devonian.

Unit 4 is part of a belt of similar Early Devonian (Upper Gedinnian)
volcanic and sedimentary rocks extending from the west end of Chaleur Bay to
the The Forks (south of Jackman) in western Maine (Boucot, and others, 1964).
From Jackman to north of Mt . Katahdin this same belt is the locus of a series
of cauldera complexes containing thick deposits of rhyolite ignimbrite dated
as Late Early Devonian (Early Emsian) . These include the Traveler Rhyolite
described by D.W. Rankin elsewhere in this Guidebook.

Unit 4 is thus part of a volcanic belt of substantial magnitude and
considerable lateral extent. The rocks are distinctly more felsic than the
earlier volcanics of Unit 1. One interpretation is that the Unit 4 volcanics
evolved on a block with at least partly-developed continental affinities, whereas
the Unit 1 volcanics were developed in a more oceanic environment.

Brown (1979) determined paleomagnetic pole positions for samples of
the Unit 4 volcanics and noted their similarity with pole positions from volcanics
of the Late Devonian Perry Formation on either side of Passamaquoddy Bay in
coastal Maine and New Brunswick. Despite the significant difference in age
between Early and Late Devonian she deduced in effect that the eastern and western
flanks of the Merimack Synclinorium were part of a coherent block* by Devonian
time. She further noted that the poles for this block differ significantly from
those of Catskill and other Late Devonian units deposited on certain North-
American basement. She argued from these data that the Presque Isle area was not
part of North American plate in Early Devonian time.

The author considers her designation of this block as Avalonian to be a serious
misnomer. None of her samples come from the Avalon Province proper and the autho:
has argued elsewhere that the term "Avalonian" should not be applied west of the
Dover (Nfld.) and Bloody Bluff - Lake Char (SE Mass. and Ct.) Faults

173

The present author notes that the youngest marine Devonian rocks
(including the Middle Devonian Mountain House Wharf Limestone of the Memphremagog
area, southern Quebec) in the Northern Appalachians occur in the Connecticut
Valley - Gaspe Synclinorium west of the Presque Isle area. It is possible that
a true ocean basin persisted there in the Late Early and Early Middle Devonian.
It is further possible that subduction of this block eastward beneath the rocks
of Unit 1 played some role in the evolution of the volcanic rocks of Unit 4.

Unit 5 contains red, arkosic sandstone and conglomerate of the Mapleton
Formation yielding plant fossils dated as late Middle Devonian (Early Givetian) .
These rocks (Stops 9 and 10) rest with angular unconformity on Unit 4 and contain
clasts of Unit 4 and older rocks. This demonstrates that an early phase of the
Acadian Orogeny occurred between Mid-Early and late-Middle Devonian time (Boucot,
and others, 1964) . The Presque Isle area is currently the only locality in the
Northern Appalachians where a significant Acadian unconformity can be bracketed
so closely.

It is noteworthy that the Mapleton Formation is itself folded, and
faulted, with dips up to 50 degress (part of which, but not all, may be original
dip of the coarser sediments). Boucot and others (1964) noted that Carboniferous
rocks at Plaster Rock, New Brunswick (about 80km east of Presque Isle) are flat-
lying and tentatively cited. This as evidence for the presence of later phases
of the Acadian Orogeny.

MISCELLANEOUS OBSERVATIONS

The work-a-day exposures in the Presque Isle area are those occurring
in the ditches along farm-roads and highways. They provide the most continuous
sections for describing the units and working out stratigraphic relationships.
Unfortunately they come and go depending on the industriousness of the road-
maintenance folk and the severity of cloudbursts. Stop 5 is an exposure of this
type. It was very well exposed in 1961 when used as the type section for what
now the New Sweden and Lower Jemptland Formations. It was only moderately
exposed when the road-log was prepared in August, 1979.

The potato economy of Aroostook County depends significantly on the
calcareous rocks of Unit 2. Where a single farm spans the boundary between
Unit 2 and Unit 3 the land underlain by Unit 2 is planted in potatoes and the
remainder is used as pasture land, chicken yards, and woodlots. Potatoes can
be grown on other units, but apparently with greater difficulty, as witnessed
by the much greater proportion of uncleared land. The field trip route crosses
this boundary on several occasions, and the participants are invited to be on
the lookout for this interesting phenomenon as a relief from the tedium of
driving.

Despite the extensive glacial transport of the regolith, it is
interesting to note that the chemistry of the soil closely reflects the under-
lying bedrock. This was noted by D. Smith, who was engaged by the Maine Geological
Survey in 1962, to reconnoiter the Unit 2 - Unit 3 contact from Washburn to
the Canadian border. He found that the cruptic soil-chemistry codes on the Soil
and Conservation Service photomosaics could be used to rough in the contact to
within an accuracy of about 1/4 mile as confirmed by his subsequent bedrock mapping.

174

REFERENCES

Boucot, A.J., Field, M.T., Fletcher, R. , Forbes, W.H., Naylor, R.S., and

Pavlides, R. , 1964, Reco-naissance Bedrock Geology of the Presque Isle
Quadrangle Maine: Maine Geol. Survey, Quadrangle Mapping Series no. 2, 123p.

Brown, L., 1979, Paleomagnetic Results from Northern Maine and the Western

Limit of "Avalon" in the Mid Paleozoic: Geophysical Res. Letters, Vol. g,
pp 821-824.

Hamilton-Smith, T. , 1970, Stratigraphy and Structure of the Ordovician and

Silurian Rocks of the Seigas Area, New Brunswick: New Brunswick Dept. of
Natural Resources, Mineral Resources Branch, Report of Investigation No. 12.

Rast, N. and Stringer, N., 1974, Recent Advances and the Interpretation of

Geological Structure of New Brunswick: Geosicence Canada, Vol. 1, pp 15-25.

Roy, D.C., and Mencher, Ely, 1976, Ordovician and Silurian Stratigraphy of

Northeastern Aroostook County, Maine: Geol. Soc. America, Memoir 148, pp 25-52

Williams, H. S. and Gregory, H. E. , 1900, Contributions to the Geology of Maine:
U. S. Geol. Survey Bull. 165, 212 p.

175

ITINERARY

ASSEMBLY POINT - HAYSTACK MT. Maine Rt 163 (Ashland Rd.) approximately 10 miles
west of Presque Isle. (Rt 163 is a left-turn off U.S. 1 about a mile
north of the college.) You will see Haystack as an obvious feature on
the western horizon shortly after leaving town. Park where the highway
crests at the southern flank of the mountain. On the left is a gravel
strip for car parking (face for departure to west) ; higher-slung vehicles
can negotiate the road that turns right from the highway crest, crossing
a small meadow to the base of the mountain.

The official trip will leave this stop at 9:00 AM. Climb as high as

ambition and time permit. If you do not have time for the climb note
road-cut exposures 50 m. east of the parking area, but watch traffic.

00.0 STOP 1 Felsic volcanic rocks of the Winterville Formation.

00.7 Castle Hill Picnic grounds on left

03.2 Left off Rt. 163 on dirt track leading into abandoned gravel pit.

(100 yards beyond turn, Rt 163 crests a low rise, drops, then climbs
steeply beyond. If you overshoot, there is a wide gravel shoulder on
right for turning about 0.1 mi beyond the low crest. Drivers of
low-slung cars may wish to turn here and park on south shoulder of
Rt. 163. Cars will depart to east.)

STOP 2 At first sight, this appears to be a gravel pit with glacial
boulders of basalt breccia, but on closer inspection one can see that
the operation has scraped down to bedrock at several places near the
back of the quarry. One exposure is of richly fossiliferous mafic tuff.
The fossils are considered Ashgillian (Late Ordovician: R.B. Newman,
oral communication, 1980) .

05.8 Castle Hill Picnic ground on right.

06.5 Rt. 163 crests at Haystack Mtn.

09.6 Left off Rt 163 on paved cutoff behind Exxon Station where Rt 163
curves to right.

09.7 Left (West) on paved road

09.9 Straight (West) at jet.

10.6 Straight (West.) Dudley Farm both sides of road. Site of early 1940's
test pits exploring low-grade manganese deposits in the New Sweden Fm
(lower Perham) and limestone-breccia "reefs" at the base of the
Jemptland Fm (upper Perham) . Note paucity of potato-fields to west

in land to west underlain by non-calcareous rocks below the Jemptland
Formation.

11.7 Left turn (North) onto Turner Road near site of former Pyle School.

176

12.3 Small gravel pit on right exposes red shale of New Sweden Fm (Lower
Perham) .

14.4 Straight (North) at stop sign crossing Maine Rt. 227

14.5 Veer left (North) on gravel road

15.8 Straight (North) Summit of Richardson Hill. Site of good section

through Castle Hill Volcanics and overlying units before "improvements"
raised road.

17.8 Park on right at bottom of hill before road curves right (East). Walk
(North) to base of field on left to Aroostook River (beware furrows and
holes hidden by over growth); west about 0.1 mile along river to
prominent outcrop.

STOP 3 The exposure displays a gentle anticline in the sandstone and

shale member of the Frenchville Formation. Note also the graded

bedding. This is the most northerly exposure of the unit on the
Castle Hill anticline.

Return to cars and continue eastbound.

20.1 Park by low cuts on right in broad right curve in road

STOP 4 The roadcut exposes micrite and slate of the Carys Mills
Formation. Further east in outcrops along the Aroostook River (the
first of which is just around the bend and down the steep bank) the
percentage of micrite is higher, and the unit has the characteristic
appearance to which the name "ribbon-rock" was given. The roadcut was
was one of the first exposures to yield fossils. They are very small
graptolites, which are best seen on slabs pulled from the cut and
washed by the rain. The fossils here are probably Ashgillian. The
discovery of younger fossils in the river bank exposure demonstrates
that the rocks locally face eastward. A fault is used to explain why
the next outcrop westward is also younger (New Sweden Formation) .

20.8 Veer left at junction with better road.

21.3 Straight at jet across bridge.

21.6 Left (Northwest) on Maine Rt. 164 into Washburn.
22.5 Left (West) on Maine Rt. 228 towards Perham.

23.8 Straight on paved secondary road as Rt . 228 curves to left.
26.5 Left (West) at junction.
27.1 Right (North) at junction.

27.4 Hedgerows on left (West) with field-road on north side. Make U-turn
leaving room for other drivers to maneuver. Park on right facing south.
Rocks for this stop exposed along field road: cars will depart to south
the way they came in.

177

STOP 5 Boucot and others (1964) designated the section along the
farm road as the type section of the Perham Formation. Roy later
elevated the Perham to Group status and named the Jemptland (poorly
exposed at the top of the hill) and New Sweden Formations (most of the
section on the east flank) . Redefinition of the gradational contact
at the base of the New Sweden Formation places the lower part of the
section (approximately 100 meters) in the Carys Mills Formation.

27.7 Right (West) at jet.

28.6 Veer left (West) at jet.

29.1 Left (southwest) on Maine Rt. 228.

Stay with Rt. 22 8 as it curves left then right through Perham.

29.4 Park on right opposite white house beyond white church at low road cut.

STOP 6 Typical silts tone of the New Sweden Formation. This is a good
outcrop for working out cleavage - bedding relationships, but be
cautious about what you call bedding here.

33.6 Stay on Rt. 228 through Washburn. Avail yourself of grocery store
here if you need "fixins" for dinner (see below: in the County,
"dinner" is the noon meal even if eaten out of a pail.)

33.8 Right (South) on Maine Rt. 164 through Washburn.

34.8 Right (West) off Rt. 164 across tracks and bridge.

34.9 Straight (West) at jet.

35.6 Curve left (South) with better road at jet.

38.5 Left (East) at jet with Maine Rt. 227.

38.8 Right (South) at jet taking road to Maple ton.

39.6 Left into Town of Maple ton Picnic Area for LUNCH STOP.
42.3 Left (East) on Maine Rt. 163 at jet in Maple ton.

43.6 Park on right at crest of hill at roadcut.

STOP 7 SPRAGUEVILLE FORMATION.

Fine calcareous siltstone characteristically showing pervasive bio-
turbation. (in good exposures the latter feature is useful in separating
this unit from the underlying Carys Mills Formation.) The rocks are cut
by teschenite ( anal cime -bearing, mafic) dikes that are better exposed
on the hill to the south.

43.9 Right off Rt. 163 (South) on Pelke Rd.

44.1 Straight. Teschenite dikes exposed in low cuts on right.

178

44.4 Veer left (South) at jet with better road.

46.1 Left (East) on road towards base of Edmunds Hill.

46.7 Left (North) at jet. Loose blocks of fossiliferous Chapman Sandstone
in rockpiles on right.

47.1 Park on right at gate to Trombley Quarry. Don't block entry road.

47.1 STOP 8 CHAPMAN SANDSTONE AND EDMUNDS HILL ANDESITE.

The best exposures are in the cleared pavement above the quarry face.
In May, 1978, the interfingering of the sandstone with the volcanics
was well-displayed, but the face may be expected to change for better
or for worse as the quarry is worked. The two units are superficially
similar (the Chapman being rich in volcanic clasts) , so the eye must
be attuned to the differences in texture between the igneous and
sedimentary units. All of the rock types may be seen in the boulder-
pile at the junction (46.7) if you cannot get into the quarry. The
quarry was opened in the late 19 70' s.

47.6 Stay with main road. Low cut exposing andesite breccia. Fossiliferous
pyroclastic rocks exposed in field reached by road to south.

48.4 Curve left (North) with main road.

49.5 Right (East) onto Maine Rt. 163.

51.1 Left (North) off Rt. 163. Park on right without blocking dump entrance.

(Departure to north) . Walk back to Rt. 163 and uphill (West) to small cut.

STOP 9 Conglomeratic facies of the Mapleton Formation. PLEASE DO NOT
HAMMER ON THIS CUT, and please refrain from prying clasts loose; this
is the only exposure the conglomerate conveniently situated for field
trips. Note the rounding of the clasts and their variety, including
both sedimentary and volcanic materials. Fossiliferous clasts of Chapman
sandstone have been reported. It is uncertain whether clasts of the
nearby Munson's Granite have been identified; a possible candidate
shown to the author by D.C. Roy about 1965 may have been a porphyritic
rhyolite.

52.7 Curve right (North) at jet.

52.8 Right (East) at crossroads.

53.1 Park on right. Walk south into woods to small quarries.

STOP 10 MAPLETON SANDSTONE plant fossils. This locality at the crest of
Winslow's Hill lies near the axis of the syncline that folds the Mapleton
formation. The woods to the south of the road contain numerous small quarry-
pits, and plant fossils are abundant at and near several of these. Plant
micro-fossils are also abundant in these rocks.

Additional exposures of Mapleton Congl. about 1 mi. further along road.

Turn back. Retrace Route to Presque Isle Dump (see 51.1)
Left (East) onto Maine Rt. 163 into Presque Isle.

179
TRIP B-8
WISCONSINAN GLACIATION OF NORTHERN AROOSTOOK COUNTY, MAINE

Andrew N Genes and William A. Newman *
Boston State College and Northeastern University, Boston

Introduction

The purpose of this trip is to examine multiple till exposures,
stratigraphic relationships, and glacially overridden deposits which permit
an interpretation of the mode of till emplacement of late Wisconsinan de-
glaciation in northern Maine. We will make a west to east transect ( Fig.
1, St. Francis to Grand Isle ) along the southern bank of the St. John
River where Wisconsinan glacial deposits are readily observable.
Maps: St. Francis ,Winterville, Eagle Lake, Fort Kent, Frenchville, Grand
Isle, Van Buren, 15 minute series.

Stratigraphy

Surficial stratigraphic units have been mapped along the Maine side
of the St. John River and yield evidence for two distinct glacial phases
represented by two tills and associated outwash. The upper, or Van Buren
Till, is the surface unit in northern Aroostook County. Thi6 till term-
inates at, and is included in, a moraine complex extending discontinuously
from St. Francis to Grand Falls, N.B. ( Fig. 1 ). Underlying the Van Buren
Till at several localities along the St. John River, the St. Francis
Till records an earlier event.

Lithology and Provenance

These two tills and their associated outwash deposits contain diff-
erent lithologies. The Van Buren drift is characterized by inclusions of
Canadian Shield Precambrian granite gneiss clasts ( varying between 2
to 5% ) and other exotics. The St. Francis drift is almost totally
supported by the local and extensive Seboomook Formation of Devonian
age ( Fig. 2 ) . No granite gneiss inclusions have been identified in the

* Work supported by the Maine Geological Survey

180

E S K ERS

STRIAE

K^RIvlere Verte
^C^s^^/p^* > Frenchvlllt^Grand Isle

Van Bureh

y younger
o older

MORAINE
) observed
inferred

1 St. FRANCIS

2 WINTERVILLE

3 POR T AGE

4 CARIBOU ~~

5 Ft. FAIRFIELD

6 STOCKHOLM

7 CASTLE HILL

8 SQUA PAN

9 MARS HILL

10 LOWER MACWAHOC

11 ORIENT

'Stockholm

\° ••••' 'VV- .;:•;.:.• .Ft. Fairfield

Castle Hill

Ashland*

■\\H\V

■ Mars
Hill

ORI FT

VAN BUREN
MARS HI LL

A

\\

Vv

10

\

\\IWK

k m

50

Fig. 1. Generalized surficial geology of Aroostook County

181

100 200

Kilometers

PROVENANCE LOCALITIES

1 Deboullie Gronodionte

2 Mopleton Sandstone

3 Chapman Sandstone

4 Mars Hill Conglomerate

\ \ M Pennsylvaman sedimentary rocks
Devonian granite plutons

Devonian sedimentary rocks includes some later Paleozoic rocks

'^=P1

IPz \A\yi pre-Devonian Paleozoic sedimentary and volcanic rocks
I Hill Precombrian granite or granite gneiss

L

PC c / / ^ Precombrian charnokite or granulits
<■ *" " Chomploin Thrust Fault

Fig. 2. Generalized geologic map of northern Maine and
adjacent Canada.

182

St. Francis drift. Erratics, glacial striations, and landforms indicate
that the Van Buren Till was emplaced by ice moving in a southeasterly
direction. Till fabric of the St. Francis Till suggests lateral east-
west movement. Located stratigraphically below and containing different
lithologies than the Van Buren Till, the St. Francis Till is not con-
sidered to have formed from the same ice regime responsible for the depos-
ition of the Van Buren Till.

Chronology and Correlations

We are primarily concerned with the question of how these two tills
correlate with other tills identified in areas adjacent to Aroostook
County and to the surface till south of the moraine complex, the Mars
Hill Till. The primary stratigraphic/chronologic difficulty which is
recognized results from the presence of granite gneiss clasts in the Van
Buren Till and their absence in both the underlying St. Francis Till and
Mars Hill Till to the south.

The difficulties in the correlation are as follows : if the Van
Buren drift was deposited by ice originating north of the St. Lawrence
River, the gneiss must have crossed the valley prior to the incursion
of the Champlain Sea at 13,000 years BP., because there is no evidence of
an ice advance through the Champlain Sea subsequent to this time. However,
it also appears that the deposition of the Van Buren Till must have occurred
after the formation of the coastal moraine belt at Pineo Ridge at about
12,700 years BP. because Aroostook County was almost certainly covered
with ice until deglaciation followed the deposition of the coastal
moraines. It would thus appear that the Van Buren Till is older than 13,000
years BP., but younger than 12,700 years BP. Even if the opening of the
Champlain Sea and the deposition of the coastal moraine belt are judged to
be synchronous, it is still difficult to explain the transportation of
the granite gneiss to the Van Buren area by an ordinary re-advance of
ice from north of the St. Lawrence Valley.

Fig. 3 shows possible correlations of Aroostook County deposits and
surrounding area deposits with the stratigraphy of southeastern Quebec.
The time constraints imposed on the correlations are those discussed

183

Strat igraphic Column, South-
eastern Quebec, (KcDonald
bnd Shilts, 1971

Stratigraphic Columns of Aroostook County, Maine, and Pertinent
surrounding areas.

Time-

StratiaraDhic

Unit

Rock-

StratigraDhic

Unit

VB
MH

VB
MH

VB

MH

VB

E VB=MH

Post

Lennoxville

Sediments

Lennoxville
Till

Van Buren
Grand Falls

/an Buren
jrand Falls

Highland F. -lighland F

St. Antonin

it. Antonin

Pineo Ridge 3ineo Ridge

Van Buren
Grand Falls

?

Mars Hill
Highland F.
St. Antonin
Pineo Ridge

Mars Hill
Highland F.
St. Antonin
Pineo Ridge

Van Buren
Mars Hill
Highland F.
Pineo Ridge

Gayhurst
Formation

Outwash

over

St. Francis

Till

Outwash

over

St. Francis

Till

Chaudiere
Till

St.FrancisTil1
'lars Hill Till

St. Francis TiWan Buren
Grand Falls

Outwash

over

St. Francis

Till

St. Francis Till

Massawippi
Formation

Outwash

over

St. Francis

Till

Outwash

over

St. Francis

Till

Johnville
Till

St. Francis
Till

St. Francis
Till

Fig. 3. Till correlation chart.

184

above and the date of 12,700 years BP. for the formation of the Highland
Front Moraine in Quebec. ( See, Introduction, Outline of the Pleistocene
geology of northern Maine and adjacent Canada, this volume ). Overrun
sediments along the southern bank of the St. John River Valley and
stratigraphic relationships indicate that the Van Buren Till is younger
than the St. Francis Till at the localities where they are juxtaposed.
If both the Van Buren Till and the Mars Hill Till are interpreted as
being approximately the same age ( Column E, Fig. 3 ) and both corr-
elative to the Lennoxville Till of Quebec, then the time restriction
imposed by the opening of the St. Lawrence Valley to the Champlain Sea
becomes a manageable problem.

Because of markedly different sediment parameters between the two
surface till units, and other considerations, the Van Buren and Mars Hill
Tills are interpreted as having been deposited penecontemporaneously as
the result of coalescing ice sheets ( Genes and Newman, 1979 ) or, as
the result of the thermal regime existing within a single ice mass ( Hughes,
pers . comm. ) .

References

Genes, A.N., and Newman, W.A., 1979, Late Wis cons inan glaciation of

Aroostook County, Maine: Geol. Soc. America Abst., Northeast Section,
p. 36.

Itinerary

Mileage

0 Assembly point for the trip is at the IGA parking lot, Fort
Kent (Junction of routes 11, 1, 161). Starting time 8: A.M.
Turn west out of parking lot on to route 161.

0.4 American customs on right

185

0.6 St. John River flood control project on right constructed by
U.S. Army Corps of Engineers.

1.5 Flood plain of the St. John River. Note terraces and incised
outwash deposits.

5.3 Crossing the St. John town line.

6.6 On left is the farm of Sylvio Martin, seed potato farmer. No
tresspassing without permission.

13.6 Outcrops of Seboomook Formation on left. This slate underlies
most of the field trip area.

18.6 Rankin Rapids campground on right.

19.9 Entering the St. Francis Moraine complex. Some bedrock is evident^
however, most of the hills are not bedrock cored.

21.9 Chamberlain general store on left. Ernie Chamberlain was intrumental
in trailblazing for the lumber companies in years past and is one
of the well known residents of Aroostook County.

22.2 Stop_l. Park along highway. Walk down path behind cabin to the

St. John River. Walk east (crossing a small stream) approximately
200 yards.

This locality ( Golden Rapids ) exhibits the entire section of
known Wisconsinan glacial stratigraphy in northern Maine. Resting
on bedrock outcrop of the Seboomook Formation is the compact, silty
to clayey, dark grey St. Francis Till, 3-3.5 meters thick. Clasts
in this till are almost wholly derived from the underlying slate.
Associated outwash, 5 meters thick, is lithologically similiar
to the till it overlies. No granite gneiss occurs in this till
or outwash. Above the outwash is the Van Buren Till which is
clayey to silty, compact, and buff to dark brown. This till and
its associated outwash contains clasts of granite gneiss and other
exotics. The St. Francis and Van Buren Tills, where they are jux-
taposed, are separated by a barely discernible boulder pavement
which consists of Seboomook Formation clasts. Near the top of the
section ( across the road from where the cars are parked ) a third
till occurs , capped by yet another thin outwash deposit. Because
of this singular occurrence and similiar lithologic constitution,
this till is assumed to be part of the Van Buren Till (ablation
facies or slight fluctuation? ) and not reflective of a separate
stratigraphic event. As slumping has occurred at various places
at this locality, care should be exercized in examining the sec-
tion.

Return to cars promptly. Reverse direction and head east.

22.4 Note outwash on left. This is the upper outwash of the strat-
igraphic section seen on stop 1. The outwash contains granite
gneiss clasts.

186

22.6 Sand dune field on left. ( Stop, time permitting ).

24. 1 §to.P 2^ Kame complex. Turn right into pit.

We are standing at the distal margin of the St. Francis Moraine
complex. This large kame, some 80' high is in contact with a thick
sequence of Van Buren Till immediately to the west. Because of
excavation, good exposures show beds dipping outward around the
entire perimeter of the deposit. The lithologies contained at
this locality are those associated with Van Buren Tills and
outwash deposits .Granite gneiss clasts in excess of half a meter
have been found at this site.

Return to cars. Turn right out of pit and continue east.

24.3 The dirt road on the right between the two green houses ( green
in 1979 ) extends through the St. Francis Moraine complex for
approximately one half mile. Pitted, hummocky, morainal topo-
graphy characterizes this region. West of this hummocky moraine
outwash extends to the St. John River and caps the Van Buren Till
seen at stop 1.

26.3 Stop 3. Turn left into Rankin Rapids campground. Drive 0.1 mile
to the large boulder at the end of the road. Go down the path to
the St. John River.

Directly at the base of the path is an exposure of the St. Francis
Till overlain by outwash of the Van Buren Till. Note the absence
of granite gneiss or " colorful " lithology of the St. Francis
Till. At times of low flow of the St. John River, it is possible
to see the St. Francis Till resting directly on the Seboomook
Formation ( approximately 50 meters to the east ). Here, till
fabrics yield a strong 90 maximum.

Return to cars and return to main road ( route 161 ) .

26.5 Proceed east on route 161.

29.1 St. Francis town line.

45.1 Continue on route 161 east towards Madawaska. Fort Kent Block-
house on left. The blockhouse marks the formal termination of
canoe trips along the famed Allagash River.

45.7 Junction. Continue on route 1 toward Madawaska.

54.7 Frenchville town line.

54.9 Turn right just before the white house with the barn which is
about to fall down.

55.1 5l2P_^i ^t£^_2H£_^2E_^rHc!S£IE^~_^--?:_5££iZ^_Ei£•

Excavation has destroyed much of this section, but relationships
are still clearly visible. Proglacial (?) Van Buren outwash

187

has been overridden resulting in recumbent folding and thrust
faulting on a large scale. Photographs of the original section
will be available if the pit is virtually destroyed by the time
we get there. This locality is directly north of the Caribou
Moraine ( to be visited on trip C - 6 ) and is thought to
represent the event responsible for the formation of the
moraine.

Return to cars and return to the main road.

55.4 Proceed south on route 1.

58.6 Turn right onto road to northern Aroostook Airport.

58.7 Turn right at sign to airport. Continue on main road.

62.4 Take left fork by stop sign. Long Lake is on the right. Continue
on main road.

64.9 Take right on Grand Isle road just before the EXXON sign in front
of the general store.

71.9 Take right at four corners. White house on hill to the right.

Across from the white house is a two striation locality ( possible
stop ) .

75.2 Turn left at main road. Entering Lille sign to the right.

75.4 §top_5^ Turn left into gravel pit. Light green house just
before entering the pit.

Here, the Van Buren Till, 3-5 meters thick, has overridden
outwash deposits associated with the St. Francis Till. The
outwash is rasped, torn, and large inclusions of the underlying
deposit is included within the overlying Van Buren Till.

Return to cars and head for home.

End of Trip

188
TRIP B-9

DEGLACIATION OF THE EDMUNDSTON AREA
AND REAPPRAISAL OF GLACIAL LAKE MADAWASKA INTERPRETATION

Claude Gauthier (Terrain Sciences, Geological Survey of
Canada) and Jacques Thibault (Dept. of Natural Resources,
New Brunswick)

The history of Glacial Lake Madawaska is closely linked
to the last deglaciation in Central and Northern New
Brunswick as well as in Northern Maine. New evidence suggest
its outlet may have been to the north, across the
Appalachian Mountains, into the St. Lawrence Valley
(Gauthier, 1980). This hypothesis will be investigated
during the comming 1980 field season and in the light of
mapping by Thibault (1979, 1980) . The main implications of a
northward outlet are: 1) persistence of ice in Central New
Brunswick and the consequently limited (if not inexistent)
influence of the Laurentide ice during deglaciation, and 2)
damming of the St. John River by ice in the Grand Falls
region and resulting in the the formation of Glacial Lake
Madawaska.

Chalmers (1885, p. 41-42 GG) described "chains of
lakes" formed in Saint John Valley behind drift dams (the
most prominent one being a frontal moraine (Lee, 1959) at
Grand Falls) . Breaching of the drift in the valley gradually
lowered the lake level and drainage was established along a
new course. Kiewiet de Jonge (1951) traced the extent of
this lake up to Lac Temiscouata, Quebec and gave it the
name: "Glacial Lake Madawaska"; he recognized only one
single phase related to the existence of the lake. Lee
(1953, 1955) confirmed this interpretation and traced the
extent of the shoreline deposits in the region of
Edmundston. Martineau (1979) observed rhythmite deposits in
the Lac Temiscouata area; he also reported the existence of
two opposite glacial flows in the area: an early flow
towards the southeast (Laurentide ice) followed by a flow
with reversed direction (Appalachian ice) . Martineau did not
present the potential regional implications of his finding.

In August 1979, during a one day field trip excursion
in the area around Edmundston, several striation sites were
observed. Convincing evidence of striated and polished
outcrops indicate that glacial flow was active towards the
west and the northwest, controled by the orientation of the
Saint John Valley and some of its tributaries. Sites were
observed along the Madawaska and the Iroquois valleys (six
sites with northwest flow) as well as along the Saint John
River near the city of Edmundston (three sites indicated a
westward flow). No other flow direction was noted, and all
striated outcrops reflected a single ice flow direction
(with maximum variations in the order of 20 degrees) . Sense

189

of flow was obtained from numerous and distinctive
characters: on a microscale, crag-and-tail and nailhead
features, and plucking of lee sides of outcrops; on a
macroscale, well developed stoss-and-lee relationships on
outcrops, producing whalebacks and roches moutonees.
Although the number of observation sites is small, it is
believed on the basis of 1) the uniformity of the various
observations, 2) the intensity of the erosional features,
and 3) the absence of other movements, that the stiae
represent the last glacial flow in the area, and that this
flow was channelled by the major valleys of the area.

Implications of the New Observation
The present observation suggests a new interpretation
of the mode of deglaciation of the Saint John Valley and of
the origin of Glacial Lake Madawaska. In order to generate a
northwestward-moving ice mass, the glacier would have had to
be centred over the highlands of central New Brunswick (and
Maine?) and it would have had to flow radially in several
directions. At one time, the frontal position of the ice was
located in the Grand Falls area, forming the so-called
Glacial Lake Madawaska by ice damming the valley.
Subsantiating evidence has been collected on the east side
of the hypothetical icemass in the Bathurst area (Gauthier,
1979 and 1978) .

The region of Plaster Rock (28 km east-southeast of
Grand Falls) presents a set of discontinuous morainic ridges
with a relief locally greater than 60 m. From limited
evidence gathered in the field (one transversal section)
glaciof luvial stratification in the moraines (beds dipping
at 5-20 degrees towards the west) indicates ice localized to
the east side of the moraines. The ice contact ridges, in
spite of their high relief (greather than 60 m) , have no
obvious regional extent; nevertheless, they seem to
represent a position of an ice cap centred in the New
Brunswick Highlands.

Field trip
The trip logs will be presented at the end of the
summer, following the summer's field work in the area.
Conclusions will be reached at that time concerning the
significance of the westward glacial flow and its regional
extension. Classical sections along the St. John River at
Grand Falls, striation sites, the Plaster rock moraines and
related observations will be included in the trip.

References
Alcock, F.J. 1948: Problems of New Brunswick geology;

Transactions of the Royal Society of Canada, v. LXII,
series III, section 4, p. 1-15.

190

Chalmers, R. 1885: Report on the Surface Geology of Western
New Brunswick; Geological Survey of Canada, Report of
Progress, 1882-1884, part XI, GG, 47 p.

Gauthier, R.C. 1978: Quelques interpretations de

l'inventaire des depots de surface, Peninsule nord-est
du Nouveau-Brunswick ; Geological Survey of Canada,
Paper 78-1A, p 409-412.

1979: Aspects of the glacial history of the

north-central highlands of New Brunswick; Geological
Survey of Canada, Paper 79-1B, p. 371-377.

1980: Existence of a Central New Brunswick ice cap based

on evidence of Northwestward moving ice in the
Edmundston area; Geological Survey of Canada, Paper
80-1A, p. 377-378.

Kiewiet de Jonge, E.J.C. 1951: Glacial water levels in the
Saint John River Valley; unpublished Ph.D. thesis,
Clark University, Worcester, Mass., 116 p.

Lee, H.A. 1953: Two types of till and other glacial problems
in the Edmundston-Grand Falls region, New Brunswick,
Quebec and Maine; unpublished Ph.D. thesis, University
of Chicago, Chicago, Illinois, 113 p.

1955: Surficial geology of Edmundston, Madawaska, and

Temiscouata counties, New Brunswick and Quebec;
Geological Survey of Canada, Paper 55-15, 13 p.

1959: Surficial geology, Grand Falls, New Brunswick (12

0/4) ; Geological Survey of Canada, Preliminary Map
24-1959

Martineau, G. 1979: Geologie des depots meubles de la region
du Lac Temiscouata; Ministere des richesses naturelles
du Quebec; DPV-618, 18 p.

Thibault, J. 1979: Granular aggregate resources of the

Grandmaison (NTS 21 N/9) and the Edmundston (NTS 21
N/8) map areas; New Brunswick Department of Natural
Resources, Mineral Resources Branch, Open File Report
No. 79-31.

Thibault, J. 1979: Granular aggregate resources of

Saint-Andre (NTS 21 0/4) and Aroostook (NTS 21 J/13) ;
New Brunswick Department of Natural Resources, Mineral
Resources Branch, (in print).

191
TRIP B-10

THE GEOLOGY AND DEFORMATION HISTORY OF THE SOUTHERN
PART OF THE MATAPEDIA ZONE AND ITS RELATIONSHIP
TO THE MIRAMICHI ZONE AND CANTERBURY BASIN

Rast, N.
University of Kentucky, Lexington, Kentucky

Lutes, G.G. and St. Peter, C.
Department of Natural Resources,
Mineral Resources Branch, New Brunswick

INTRODUCTION

Western New Brunswick is underlain by parts of two
tectonostratigraphic zones: the Matapedia Zone and the Miramichi
Zone. Part of the Miramichi Zone is downfaulted forming the
Canterbury Basin in the southern part of the area (Figure 1) . The
purpose of this trip is to examine sections from Florenceville to
Canterbury and contrast the structural development of the two
zones with the Canterbury Basin.

The Matapedia Zone has been mapped in the Woodstock
area by Anderson (1968) , and Hamilton-Smith (1972) . The regional
geology in the southern part of the Matapedia Zone has been
demonstrated by Pavlides (1968) . Reconnaissance work by us has
modified the structural interpretation for this area and an early
(Pre-Acadian) period of recumbent folding is recognized. Recent
mapping by the Department of Natural Resources has modified
previous interpretations of the stratigraphy, structure and meta-
morphism of the Miramichi Zone. The rocks in the Canterbury Basin
have been mapped by Venugopal (1978, 1979) and Lutes (1979) and
the general stratigraphy and structure are well known.

I,

— iV—

CANTERBURY BASIN

Conglomerate, slate, limestone, and
minor volcanic rocks

MATAPEDIA ZONE

Smyrna Mills Formation

v^V\] Grey slate, silts tone; minor greywacke,

\\\>^j calcareous sandstone and manganiferous
slate (includes Wapske Formation silt-
stone, sandstone and volcanic rocks east
of Florenceville

Caryt Mills Formation

Calcareous slate, calcareous siltstone,

argillaceous and arenaceous limestone

192

MIRAMICHI ZONE
Tetagouche Group
TTTj Quartzite, slate; minor volcanic and
'.\\- volcanoclastic rocks

/?» Trace of Ff -recumbent fold
^V^v Plunging F2 - sync I me, anticline

KILOMETERS
0

Figure I. Generalized Geology of Western New Brunswick.

193

MATAPEDIA ZONE

Stratigraphy

The Matapedia Zone strikes from northeastern Maine
across northwestern New Brunswick into the southern part of the
Gaspe Peninsula. The zone is composed of limestones, greywackes,
siltstones and shales ranging from Middle Ordovician to Late
Silurian in age. The rocks in northern New Brunswick have been
separated into four rock-stratigraphic units: Grog Brook Group,
Matapedia Group, Upsalquitch Formation, and Perham Formation
(St. Peter, 1978) . In the Woodstock area, the Matapedia Zone is
underlain by calcareous slates of the Carys Mills Formation which
are stratigraphic equivalents of the Matapedia Group and range in
age from Middle Ordovician to Lower Silurian (Pavlides, 1968) .
Its thickness has been estimated to range from 1500 to 12,000 feet
in Maine; the formation thins to the west (Pavlides, 1968) . In
New Brunswick, it has been estimated to be greater than 2500 feet
thick (Hamilton-Smith, 1972) and has been interpreted as a tur-
biditic sequence of calcareous flysch (Pavlides, 1968). The
Smyrna Mills Formation conformably overlies the Carys Mills For-
mation in Maine (Pavlides, 1968) . It is composed of slate,
calcareous slate and sandstone with minor manganif erous siltstone.
This formation ranges in age from early Llandovery to early
Ludlow (Pavlides and Berry, 1966). Parts of the Wapske Formation
occur above the Smyrna Mills Formation southeast of Florenceville .
The Wapske Formation consists of intercalated clastic sedimentary
and mafic volcanic rocks which are interpreted by St. Peter (1978)
as an alternating sequence of marine and terrestrial rocks, the
marine rocks probably being shelf deposits. The age of the Wapske
Formation is established as Lower Devonian, probably Helderbergian
(St. Peter, 1978) .

Structure

Two periods of folding have been recognized in the
Woodstock-Florenceville area. The first has produced large
recumbent nappes which are interpreted from flat-lying overturned
beds near Woodstock. There is no cleavage associated with these
Fi-folds and the age of deformation is difficult to assess. The
Smyrna Mills Formation has been demonstrated to conformably over-
lie the Carys Mills Formation near Houlton, Maine (Pavlides, 1968)
Near Woodstock (Stop 7) , the contact between the two formations
appears to be conformable, suggesting at least part of the Smyrna
Mills Formation was involved in the recumbent folding. As noted
previously, the youngest age known for the Smyrna Mills Formation
is Ludlow (Pavlides and Berry, 1966) . The oldest age for the

194

overlying Wapske Formation is Lower Devonian (Helderbergian) .
The intervening and unrecorded stratigraphic interval may repre-*
sent a period during which uplift in the adjacent Miramichi Zone
created unstable slope conditions in part of the Matapedia Basin
initiating gravity slides prior to the main Acadian Orogeny. It
is also possible that the Smyrna Mills Formation, where observed,
is paraconformable over the Carys Mills Formation. If so, gravity
sliding may have been initiated by Taconian movements in the
Miramichi Zone in Upper Ordovician time.

A superimposed Acadian generation of folding and
cleavage affects all rocks in the Matapedia Zone. The style and
attitude of these folds varies with change in orientation of the
earlier Fi-folds. On the limbs of F]_-folds, Acadian folds are
open to close and have associated axial planar cleavage. Plunge
of these folds is dependent on the attitude of the Fi-fold limb.
In the vicinity of F^-fold hinges, Acadian folds are tight and
have large variations in plunge and bedding-cleavage intersection
lineation.

MIRAMICHI ZONE

Stratigraphy

Rocks of the Tetagouche Group underlie most of the
Miramichi Zone. The rocks range from Cambrian (?) to Middle
Ordovician and can be divided into five map-units consisting of a
thick basal unit of quartzite and slated) overlain in succession
by (2) slate, siltstone and greywacke, (3)rhyolitic volcanics,
(4)manganiferous slate, chert and andesitic volcanics, and (5)
massive basic volcanics (Helmstaedt, 1971) . In southwestern
New Brunswick, lithostratigraphic equivalents of the first four
units have been recognized (Venugopal, 1979; Lutes, Unpubl. Msc).
Manganif erous slate forms a readily recognizable marker horizon
throughout the Miramichi Zone.

Structure

Rocks of the Miramichi Zone are structurally more
complex than those of the Matapedia Zone or Canterbury Basin.
First folds, which are rarely observed, have associated cleavage
parallel to bedding. Later, second folds of bedding and the first
cleavage in the Canterbury area pre-date the main third defor-
mation which folds and cleaves both Cambro-Ordovician and Siluro-

195

Devonian rocks (Lutes, 1979) . These third folds are generally
steeply dipping with axial planar cleavage. Much of the Miramichi
Zone is occupied by granitoid plutons, generally of Devonian age.
Whereas regional metamorphism has little affected rocks of the
Matapedia Zone and the Canterbury Basin, the grade of metamor-
phism in the Miramichi Zone reaches sillimanite grade within a
megmatite complex that occurs in central New Brunswick. Deforma-
tion and regional metamorphism of this zone is attributed to the
Taconian Orogeny of Middle-Upper Ordovician age and spans the time
during which the Carys Mills Formation was being deposited in the
Matapedia Basin.

CANTERBURY BASIN

Stratigraphy

Rocks of the Canterbury Basin comprise conglomerates,
slates, mafic and felsic volcanics and limestone unconformably
overlying rocks of the Miramichi Zone. These Siluro-Devonian
rocks occupy a downfaulted basin bound by the Meductic Fault on
the west and the Charlie Lake Fault on the east (Figure 1) .
Pocowogamis Conglomerate forms the base of the succession along
the east side of the Meductic Fault. This is overlain by Scott
Siding Slate succeeded by feldspar-quartz crystal tuff and mafic
volcanics. In the southern part of the area, Cambro-Ordovician
quartzite of the Miramichi Zone is unconformably overlain by
Canterbury Limestone with some basal conglomerate. The Canterbury
Limestone and Scott Siding Slate are gradationally overlain by
calcareous sandstones and siltstones of the Hartin Formation
containing intercalated mafic and felsic volcanics, microconglom-
erate and minor limestone. Fossils from the Hartin Formation
have yielded a Lower Devonian (Helderbergian) age. Underlying
formations may be Silurian.

Structure

Rocks in the Canterbury Basin occupy a downfaulted
basin (graben structure) and subsequent deformation produced north
facing folds in the south and south facing folds in the north.
First schistosity in these rocks has a similar trend and attitude
as second schistosity in Cambro-Ordovician rocks and is probably
the same age. Metamorphism by the nearby Pokiok Batholith has
formed biotite in many of the Siluro-Devonian rocks on the north.

196

REFERENCES

ANDERSON, F.D. 1968. Woodstock, Millville and Coldstream map-area,
Carleton and York Counties, New Brunswick. Geological Survey
of Canada, Memoir 353.

HAMILTON- SMITH, T. 1972. Stratigraphy and structure of Silurian
rocks of the McKenzie Corner area, New Brunswick. New
Brunswick Department of Natural Resources, Mineral
Development Branch, Report of Investigations No. 15.

HELMSTAEDT, H. 1971. Structural geology of Portage Lakes area,
Bathurst-Newcastle District, New Brunswick. Geological
Survey of Canada, Paper 70-2 8.

LUTES, G. 1979. Geology of Fosterville-North and Eel Lakes,

map-area H-2 3 (parts of 21G/13 and 21G/14) . New Brunswick
Department of Natural Resources, Mineral Resources Branch,
Map Report 79-3.

LUTES, G. 1980. The Geology of the Northwestern Part of the

Metamorphic Aureole of the Poliok Batholith. Unpublished
MSc thesis (in progress) .

PAVLIDES, L. 1968. Stratigraphic and Facies Relationships of the
Carys Mills Formation of Ordovician and Silurian Age North-
east Maine. U.S. Geological Survey, bulletin 1264.

PAVLIDES, L. and BERRY, W.B.N. 1966. Graptolite-bearing Silurian
rocks of the Houlton-Smyrna Mills area, Aroostook County,
Maine. U.S. Geological Survey, professional paper 550-B,
p. B51-B61.

POOLE, W.H. 1976. Plate Tectonic Evolution of the Canadian

Appalachian Region. Geological Survey of Canada, paper 76-1B.

ST. PETER, C. 1978. Geology of Parts of Restigouche, Victoria
and Madawaska Counties, Northwestern New Brunswick, N.T.S
21N/8, 21N/9, 210/5,210/11,210/14. Mineral Resources Branch,
Dept. Nat. Res., report of Investigation 17.

ST. PETER, C. 1978. Geology of Head of Wapske River map-area J-13
(21J/14) . New Brunswick Department of Natural Resources,
Mineral Resources Branch, Map Report 78-1.

197

VENUGOPAL, D.V. 1978. Geology of Benton-Kirkland-Upper Eel River
Bend, Map-area G-22 (21G/13E) . New Brunswick Department of
Natural Resources, Mineral Resources Branch, Map Report 78-3

VENUGOPAL, D.V. 1979. Geology of Debec Junction-Gibson,

Millstream-Temperancevale, Meductic Region, map-areas G-21,
H-21, 1-21, and H-22 (Parts of 21J/3, 21J/4, 21G/13, 21G/14)
New Brunswick Department of Natural Resources, Mineral
Resources Branch, Map Report 79-5.

ITINERARY

Assembly point is in Florenceville at stop 1, a deep
road cut on route 103 about 0.5 km southwest of the intersection
with route 2 on the west side of the bridge to East Florenceville
across the Saint John River. Park vehicles along side of road.
Assembly time is 9:00 a.m.

Kilometers Miles

00.0 00.0 Stop 1. Carys Mills Formation. The beds are

closely folded (of Acadian generation) and
plunge approximately 55° to 016. Grading and
small-scale load structures suggest bedding
is right way up and faces northward. Numerous
faults cut the section and gabbroic dykes are
common .

00.5 00.3 Intersection with rts . 2 at bridge. Proceed

across bridge and travel southward on rte. 2.

06.1 02.4 Stop 2. Small exposure on east side of road.

Carys Mills Fm. Thinly bedded and laminated
calcareous siltstones and slate. Grading
appears to be right way up on cleavage
trending 014 and dipping vertically. Bedding-
cleavage intersection plunges about 60° to
014. Structural style appears to be consis-
tent to this point. Over the next 6-7 km we
will be crossing a section of Smyrna Mills Fm.
and Wapske Fm. which overlie the Carys Mills
Fm. and appear to plunge northwesterly.

2 0.1 12.6 Stop 3. Carys Mills Fm. Current laminated,

thinly bedded calcareous sandstone and
lithographic limestone. Bedding consistently
trends 020, dips 50° to the west and is right
way up. Bedding-cleavage intersection is
subhorizontal and indicates a change in the
plunge of Acadian folds. As the Acadian folds
are superimposed on an earlier (Fj) generation
of recumbent folds, this must reflect a change
in the attitude of first structure. We are
interpreted to be on the western limb of an
overturned recumbent Fj-fold (Figure 2a,b).

198

0. Florenceville

b. Woodstock i ^ /

B

M-f

/

/

/
€)

/

f

limestoneQ conglomerate F

C. Canterbury

FIGURE 2.
Schematic cross-sections through parts of western
New Brunswick. Trip stops are indicated.

199

Kilometers Miles

31.5 19.7 Stop 4. Carys Mills Fin. Calcareous slate and

siltstone. Small isoclinal F2 (Acadian) folds
with axial planar cleavage trending parallel
to road (024) . The plunge of these folds is
extremely variable within the outcrop as is
indicated by bedding-cleavage intersections.
This appears to be the affect of superimposed
folding over Fi-folds. We are interpreted to
be in the hinge area of a major Fi-fold
(Figure 2a, b) . Younging indicators are
inconclusive, but exposure 1 km to the south
appears to be downward facing.

32.6 20.4 F2~folds in this exposure plunge about 20° to

2 04 and bedding appears to be downward facing.

38.5 24.1 Stop 5. Intersection of rte. 2 and hwy. 550.

Carys Mills Fm. Calcareous slate and silt-
stone. Bedding is flat lying, downward facing
and openly folded by upright, horizontal F2-
folds (Figure 2b) . Grading and dewatering
structures give facing direction. We are
interpreted to be on the overturned limb of
the F^-recumbent fold.

41.3 25.8 Intersection of rte. 2 and rte. 95 to Houlton.

Turn right.

41.5 25.9 Stop 6. Carys Mills Fm. Similar lithology

and structure as stop 5. This exposure con-
tains the best criteria for downward facing.
Grading is well developed and a sedimentary
cut-off provides convincing evidence of way up,

42.3

26.4 Intersection, bear right.

143.6 17.0 Stop 7. Carys Mills Fm. is steeply dipping,
tightly folded and faces southward. An

apparently conformable contact with Smyrna
Mills Fm. is exposed to the west, on the south
side of the highway. Smyrna Mills Fm. here
consists of steeply dipping thinly bedded
sandstone and slate which also appear to be
tightly folded by steeply plunging F2~folds
which face to the southwest. We are inter-
preted to be in the hinge area of the Fl~
recumbent fold.

200

Kilometers Miles

4 8.2 30.1 Stop 8. Smyrna Mills Fm. Fine-grained green

calcarenite and slate. Large upright syncline
has very shallow plunge. Graptolites can be
found in this exposure.

49.2 30.6 Intersection with rte. 540. Turn right and

follow road past Belleville and Jackson Falls
to Oakville.

61.7 38.6 Stop 9. Oakville. Turn right, cross bridge

and park on east side of river. Polydeformed
Carys Mills consists of argillaceous limestone
with thinly bedded and laminated sandy lime-
stone. Steep F^-folds are refolded by super-
imposed F2~folds and cleavage trending 010
(Figure 2b) .

82.1 51.3 Turn and proceed back to Intersection of rte.

2 and rte. 95 at Woodstock, past stop 6. Turn
right and travel southward.

84.3 52.7 Stop 10 (optional). Strongly sheared quartzite,

quartz wacke and slate of the Miramichi Zone
southeast of the Woodstock Fault. This fault
separates the Matapedia and Miramichi Zones
and has a pronounced topographic expression
trending in a northeasterly direction.

104.1 65.1 Stop 11. Polydeformed quartzite and slate

just east of Meductic across Eel River. This
is the typical lithology of the basal Tetagouche
Gp. of northern New Brunswick and is similar
to Grand Pitch type lithology in Maine.
Small-scale, vertically plunging, asymmetric
F2~folds with axial planar fracture cleavage
fold bedding, first cleavage (which is sub-
parallel to bedding) , and L^-intersection
lineation which, if axial planar, suggests
F^-folds were originally plunging southwards.

105.1 65.7 Stop 12. Long road-cut of steeply dipping

reddish, manganiferous slate, chert and
volcanoclastic sedimentary rocks. These
overlie the quartzite-slate unit and are
interpreted to occupy the axial zone of a
steeply plunging, southward facing syncline

Kilometers Miles

201

(Figure 2c) . More quartzite outcrops to the
east, and the northeasterly trending Meductic
Fault truncates the Cambro-Ordovician sequence
approximately 2 km to the east and to the
south.

105.2

65.7

Turn vehicles and proceed southward on rte
122 towards Canterbury.

108.4

109.4

110.4

67.8

69.0

Intersection. Turn left toward Johnson Sett.

68.4 Stop 13. Small outcrop on left of strongly
cleaved polymict conglomerate. Conglomerate
is extensive in the Siluro-Devonian sequence
of the Canterbury Basin along the Meductic
Fault. Venugopal (19 79) has envisaged for-
mation of the conglomerates along an escarp-
ment formed by the Meductic Fault. Grey Scott
Siding Slate, a dominant lithology in this
area, becomes calcareous to the south and
grades into the Canterbury Limestone.

Return to main road, turn left and proceed
to Canterbury.

119.4

74.6

Canterbury. Bear right past school and turn
right past cemetery.

122.4

76.5

Intersection. Turn left and proceed south-
ward across railway tracks.

128.4

80.3

Stop 14. Pass church on right side of road
and park in drive of deserted house at bottom
of hill on left. Exposure is scattered in
back field. The rock is essentially a quartz-
ite-pebble conglomerate with a limey matrix.
This represents the base of the Canterbury
Limestone and unconf ormably overlies Cambro-
Ordovician quartzite and slate immediately to
the west from which the pebbles were derived.
The relatively higher topography marks areas
underlain by quartzitic rocks.

END OF TRIP

Return to cars and retrace to route 2.

202

TRIP C-l

STRATIGRAPHY AND SEDIMENTOLOGY OF THE

SIEGAS FORMATION (EARLY LLANDOVERY) OF

NORTHWESTERN NEW BRUNSWICK

Terence Hamilton-Smith
Sohio Petroleum Company, San Francisco

Introduction

The Siegas area extends across the international border between the
United States and Canada at the intersection of the Grand Isle, Stockholm and
Van Buren 7 1/2 minute quadrangle maps of the U.S. Geological Survey.
Geological work in the area was done between 1967 and 1969 as a Master of
Science thesis (Hamilton-Smith, 1969) at the Massachusetts Institute of
Technology. Acknowledgements are due to R. R. Shrock, E. Mencher, A. T.
Boucot, W. B. N. Berry, J. M. Berdan, R. B. Neuman, J. W. Huddle, D. C. Roy,
T. B. Griswold and G. Planansky for assistance, encouragement and guidance
during the work. Financial assistance was received from the National Science
Foundation and Department of Natural Resources of the Province of New
Brunswick .

Rocks of the Siegas area are entirely sedimentary and are between
Middle Ordovician and Late Silurian in age. They occur in the tightly folded
common limb of the Ashland Synclinorium and the Pennington Mtn . Anticli-
norium. The earliest tectonic event recorded in the rocks was the Taconic
orogeny, which resulted in the deposition of the Siegas Formation and
possible low intensity folding in some older rocks. Bulk deformation of the
area occurred after the Late Silurian, probably during the Acadian orogeny.

A simplified geological map of the Siegas area is presented in Figure
1. A homogeneous fold system is recognized which consists of plane, cy-
lindrical, tightly appressed similar folds which are slightly inclined and
plunging. Details of outcrop distribution, structural attitudes, etc. may
be found in Hamilton-Smith (1970).

Stratigraphy

The oldest unit exposed in the Siegas area is the Madawaska Lake
Formation (Roy et al . , 1976) of Late Middle to Late Ordovician in age. The
unit is more than 1950 feet thick and is composed of dark grey slate with
minor quartzose sandstone. The base of this unit is not exposed. The
sandstone is light grey, highly calcareous (25%), quartzose (70%) and fine-
grained, with minor amounts of plagioclase, biotite, lithic fragments and
sphene. This lithotype occurs in beds from 1/2 inch to 4 feet thick and is
usually laminated and cross-laminated. Graptolites collected from one
locality in the Siegas area indicate a probable Zone 13 of the Caradoc age
(Hamilton-Smith, 1970).

203

The Carys Mills Formation is between Late or Late Middle Ordovician and
earliest Silurian in age and conformably overlies the Madawaska Lake
Formation in the Siegas area. The Carys Mills Formation is about 1300 feet
thick and is composed of interbedded slate, sandstone and limestone. The
characteristic lithotype is dark grey, argillaceous calcilutite (75-90%
calcite) which weathers either light bluish grey or light brown. This occurs
in beds from 1/2 inch to 16 inches thick and is commonly crosslaminated as
modified by convolute lamination. Graded bedding is locally common but
obscurely developed, particularly in the finer grained beds. Conodonts and
graptolites collected from two localities in the Siegas area indicate a Late
Ordovician to Early Silurian (early or middle Llandovery) age (Hamilton-
Smith, 1970).

The Siegas Formation of Early Silurian (early Llandovery) age overlies
the Carys Mills Formation in the Siegas area with general conformity. Local
unconformity due to submarine erosion is established at two localities. The
thickness of the Siegas Formation ranges from 800 to 350 feet in the area.
The unit is composed of sandstone with minor conglomerate, slate and
limestone. Details of facies and lithotypes are discussed below. The
sandstone occurs in beds from 1 to 314 inches thick with a variety of
sedimentary structures suggestive of depostion from turbidity currents.
Modal compositions range from quartz arenites through arkose to lithic
wacke . The conglomerate occurs in beds from 2 inches to 27.5 feet thick and
consists of limestone clasts up to 36 inches in maximum dimension contained
in a lithic wacke matrix. An extensive collection of brachiopods from the
Siegas Quarry (EM558) indicates an Early Llandovery age (Ayrton et al . ,
1969).

The New Sweden Formation (Roy et al . , 1976) conformably overlies the
Siegas Formation and is between Early Llandovery and Ludlow in age. No
fossils are known from this unit in the Siegas area. The New Sweden Formation
is about 600 feet thick and consist of grey calcareous slate with minor
mamgani f er ous silt s t one .

The Jemtland Formation (Roy et al . , 1976) conformably overlies the
New Sweden Formation. The unit may be from late Wenlock through early
Ludlow in age but in the St. John River valley only Early Ludlow fossil
collections are known. No fossils are known from this unit in the Siegas
area. The Jemtland Formation is composed of calcareous shale, slate and
siltstone with minor sandstone and limestone and is probably more than 2000
feet thick. The top of this unit is not exposed.

Siegas Formation

The Siegas Formation is of restricted areal extent. To the southeast the
unit disappears within 8 miles, probably passing laterally into equivalent
beds of the Carys Mills Formation of Early Llandovery age. To the southwest
the unit may pass laterally into equivalent beds of the basal Frenchville
Formation. Reconnaissance mapping suggests that the Siegas Formation extends
at least 7 miles to the north and northwest and possibly as much as 40 miles
to the northeast.

204

Figure 1: Generalized geological map of the Siegas area,
New Brunswick and Maine.

205

x » » » J - - '. -r', - - '." - ', - s
^\ xN x\ »\ x\ ^»\

' ' ' *i^' * '

,»*•»*, »v,^*,v'

Jemtland Formation: thinly interbedded , fissile, laminated
shale, slate and siltstone with minor fine grained light
grey sandstone and limestone.

"'"//A////""'////

New Sweden Formation: dark to medium grey, laminated, calca-
reous slate with laminated manganiferous iron-rich siltstone

rr^^^^rrrr^'^^'Tr'TTTTiT^^'-

Siegas Formation: interbedded sandstone, slate and conglomerate
with minor limestone and chert. Sandstone composition varies
from lithic wacke to quartz arenite.

Carys Mills Formation: thinly interbedded dark grey, calcareous
slate and limestone. Limestone is dense and micritic and weathers
characteristic blue grey and light brown colors.

Madawaska Lake Formation: dark grey, sparsely laminated, non-
calcareous slate with minor light grey calcareous fine grained
sandstone .

0

Formation contact
Fossil locality

Fault
Scheduled stop

206

Within the Siegas area there are significant facies variations within the
Siegas Formation. The unit in the western part of the area is 600 to 800 feet
thick, characterized mainly by lithic wacke and limestone conglomerate. To
the southeast the unit consists of 500 to 600 feet of section, mainly
sandstones with a composition ranging from feldspathic arenite to arkosic
wacke. To the northeast the Siegas Formation is 350 to 500 feet thick and is
composed mainly of quartz arenite, slate and limestone.

The lithic wacke facies displays significant changes from south to north.
There is an increase in total section thickness from 500 to 800 feet, an
increase in the thickness and abundance of limestone conglomerate beds and the
progressive development of an erosional surface between the Siegas Formation
and the Carys Mills Formation. As much as 150 feet of the Carys Mills
Formation has been eroded in the north part of the facies, at the Siegas Quarry
section. The well exposed section in the Siegas Quarry is discussed in detail
below.

Facies variations within the Siegas Formation are interpreted as the
result of deposition by different processes in a region of complex topography.
The lithic wacke facies is understood as the product of a submarine channel-
fan system. Paleocurrent information from the Siegas Quarry indicates a
derivation from the north. The quartz arenite facies is interpreted as a
winnowed shelf deposit resulting mainly from wave action. The arkosic facies
is a finer grained deeper water equivalent of the quartz arenite facies
deposited on the slope between the shelf and the submarine fan (Hamilton-
Smith, 1971a).

The composition of sandstones of the Siegas Formation indicates a
specific provenance. Sandstones of the lithic wacke facies consist mainly of
andesite fragments and their disintegration products: plagioclase (An32
average), quartz and pyroxene. Pyroxene and plagioclase compositions of the
sandstones are identical to those of the phenocrysts of the andesite
fragments. A distinct minor petrological assemblege (up to 20%) consists of
felsic plutonic fragments and their disintegration products: potassium
feldspar and quartz. Quartz grains from these two assembleges are mor-
phologically distinct.

Sandstones of the arkosic facies consist mainly of felsic plutonic
fragments, potassium feldspar and quartz. The feldspars of the plutonic
fragments include orthoclase, perthite, sodic plagioclase and myrmekite and
suggest an original source material with a composition between diorite and
granite. A small amount of andesite fragments (up to 2%) occurs within these
sandstones .

The quartz arenites of the Siegas Formation consist mainly (87 to 95%)
of medium grained rounded quartz suggesting a polycyclic history withat
least partial derivation from older quartzose sandstones. However, the
compositions of the feldspars and lithic fragments of the quartz arenites are
identical to those of the arkosic facies, suggesting at least in part a
common provenance of felsic plutonic rocks.

207

The source area of the Siegas Formation is interpreted as a complex
local uplift in northwestern New Brunswick consisting of quartzose sand-
stone, andesite and quartz diorite (Hamilton-Smith, 1971a). This source
area would be part of the northern end of the land mass Taconica of Roy (this
volume). An analogous association of all three lithotypes is known
regionally in the WeeksboroLunksoos Lake anticlinorium (Neuman, 1967).

Siegas Quarry Section

The Siegas Quarry provides an almost completely exposed section
through the Siegas Formation in the thickest and most proximal part of the
lithic wacke facies. Each bed in the section was numbered, individually
described and assigned to one of the lithotypes: limestone conglomerate,
sandstone, siltstone, limestone or chert. A first-order Markov process
transition matrix was derived in order to define lithological associations.
The matrix elements and other features of the five lithotypes are summarized
in Table 1.

Table 1: Lithological features of the Siegas Quarry section (after Hami-
lton-Smith, 1971b).

Lithotype Transition Matrix

SI S2 S3 S4 S5

Conglomerate Si .27 .68 .05 0 0

Sandstone

S2

.07

.45

.33

.09

.06

Siltstone

S3

.01

.29

.32

.37

.01

Limestone

S4

.01

.15

.68

.15

.01

Chert

S5

0

.75

.17

.08

0

Percent

Percent

of Thickne

ss

of Beds

13

4

72

34

12

39

2

21

1

2

Inspection of the matrix suggests two empirical lithological associ-
ations: an exogenic group consisting of limestone conglomerate, sandstone
and chert and an endogenic group consisting of limestone and siltstone. In
the section at the quarry the lithotypes of the exogenic group occur in
three distinct thickly bedded intervals separated by two thinly bedded
intervals consisting of lithotypes of the endogenic group.

The limestone lithotype at the Siegas Quarry is an impure micrite
identical to the limestone of the underlying Carys Mills Formation. If a
transition matrix is derived for the endogenic group and compared to one
from the Carys Mills Formation the results are very similar (Table 2).

Table 2: Transition matrices for the endogenic group, Siegas Quarry section
and the Carys Mills Formation (after Hamilton-Smith, 1971b).

Lithotype Transition Matrix Transition Matrix

Carys Mills Formation
SI S2

.15 .85
.55 .45

Endegenic Group

SI S2

Limestone Si

.23 .77

Siltstone S2

.42 .58

208

The interpretation of the transition matrices is that the endogenic
group of lithotypes records an extension of the sedimentary conditions of
the Carys Mills Formation throughout the time of deposition of the Siegas
Formation. The exogenic group records three distinct interruptions of this
relatively continuous sedimentation by the deposition of sandstone and
limestone conglomerate.

Most of the sandstone beds display the classic characteristic
features of turbidite deposition. These features can be well observed on
the Siegas Quarry and have been fully described previously (Hamilton-
Smith, 1970, 1971b). The distribution of sedimentary structures in the
section, particularly in the most significant middle exogenic event,
indicates a passage from relatively distal to highly proximal and back to
distal depositional conditions. An important feature of the Siegas Quarry
section is the presence of thick nongraded sandstone beds and subintervals
near the base of graded beds associated with the most proximal phase of
turbidite deposition. These nongraded beds and intervals contain scattered
limestone clasts up to 24 inches in maximum dimension and are interpreted
as grain flow deposits. .

The limestone conglomerate beds are associated with the turbidite
deposition. They generally consist of limestone and slate clasts up to 24
inches in maximum dimension (60% average) and a sandstone matrix com-
positionally identical to the lithic wacke of the Siegas Quarry. The
limestone and slate clasts are lithologically identical to the lithotypes
of the endogenic group and the Carys Mills Formation. The angularity,
pull-apart structures and plastic deformation of individual limestone
clasts suggest a short history of transportation and semi-consolidation at
the time of erosion.

Half of the limestone conglomerate beds are strongly graded and have
load casts, erosional surfaces and clast imbrication at the bases of the
beds. They occur at the most proximal phase of the middle exogenic event
and are interpreted as the product of deposition from turbulent flows with
unusually high values of effective density and turbulent intensity. Such
values may have been produced by lateral confinement of the flow by a
narrow channel.

The other limestone conglomerate beds are not graded and have an
internally heterogereous structure. They were probably produced by var-
ious forms of mass flow from incoherent slumping to a complex sliding of
relatively coherent lenticules over one another with most of the shear
stress concentrated rear lenticule margins.

The limestone conglomerate beds orginated from a mixture of lime-
stone and slate clasts and unconcolidated sand consisting mainly of
andesite fragements. The limestone and slate was probably derived from
submarine erosion of the Carys Mills Formation a short distance to the
north of the Siegas Quarry. Mixing with the sand may well have occurred
by collapse of submarine canyon walls into unconsolidated sand forming the
channel floor. Such events may also have initiated the high density for
trubulent flows that resulted in the deposition of graded limestone
conglomerate beds .

209

References

Ayrton, W. G. , W. B. N. Berry, A. J. Boucot , J. Lajoie, P. J. Lesperance,

L. Pavlides, and W. B. Skidmore, 1969, Lower Llandovery of the north-
ern Appalachians and adjacent regions: Geol . Soc . Amer. Bull., v. 80,
p. 459-484.

Hamilton-Smith, T., 1969, Sedimentation during the Taconic orogeny: a study
of late Ordovician and early Silurian rocks of the Siegas area, New
Brunswick: S. M. and S. B. thesis, Massachusetts Institute of Tech-
nology, Cambridge, Massachusetts.

Hamilton-Smith, T., 1970, Stratigraphy and structure of Ordovician and Sil-
urian rocks of the Siegas area, New Brunswick: New Brunswick Dept .
Natural Resources, Mineral Resources Branch, Rept . Investigation No.
12.

Hamilton-Smith, T., 1971a, Paleogeography of northwestern New Brunswick

during the Llandovery, a study of the provenance of the Siegas For-
mation, Can. Jour. Earth Sciences, v. 8, p. 196-203.

Hamilton-Smith, T., 1971b, A proximal distal turbidite sequence and a pro-
bable submarine canyon in the Siegas Formation (Early Llandovery) of
northwestern New Brunswick: Jour. Sed . Pet., v. 41, no. 3, p. 752-
762.

Neuman, R. B., 1967, Bedrock geology of the Shin Pond and Stacyville qu-
adrangles, Penebscot County, Maine: U.S. Geol. Survey Prof. Paper
542-1.

Roy, D. C. and Ely Mencher, 1976, Ordovician and Silurian stratigraphy

of northeastern Aroostook County, Maine: Geol. Soc. Amer. Mem. 148,
p. 25-52

Itinerary

Mileage

0 Assembly for trip in Presque Isle, Maine. Starting time 7:45 a.m
Drive north on Route 1 to Van Buren, Maine.

33.0 Enter Van Buren, Maine.

34.0 Cross St. John River into St. Leonard, New Brunswick. Positive

identification of citizenship is required for crossing and return.
Drive northeast on Route 17 to intersection with TransCanada High-
way .

35.3 Turn left onto TransCanada Highway and proceed westbound towards
Edmunds ton.

210

35.9 Stop 1 . Park along shoulder of the road well clear of the highway.
Outcrop is in the road cut on both sides of the highway.

New Sweden Formation. The exposure is located on the eastern limb
of the Ashland Synclinorium and consists of tightly folded light
grey, calcereous laminated slate with laminated manganiferous and
iron-rich siltstones.

Return to cars. Proceed west on highway.

37.8 Stop 2. Park along shoulder of the road well clear of the highway.
Outcrop is in the road cut on both sides of the highway.

Jemtland Formation. The outcrop is to the west of the axis of the
Ashland Synclinorium just east of the map-area of Figure 1. The
outcrop consists of thinly laminated, fissile, calcareous siltstone,
shale and sandstone.

Return to cars. Proceed west on highway.

4 0.5 Stop 3. Park along shoulder of the highway well clear of the road.
Outcrop is in the road cut on the north side of the highway.

Carys Mills Formation. The exposure consists of steeply dipping

interbedded dark grey calcareous slate and limestone with minor

snadstone. The limestone beds weather light blue grey or light

brown. The sequence occurs near the base of the Carys Mills For-
mation.

Walk west along the highway to Stop 4.

Stop 4. The outcrop is in the road cut on the south side of the
highway. Cross with due negard for the occasionally very fast
traffic .

Carys Mills Formation and Siegas Formation. From east to west the
outcrop consists of thinly interbedded dark grey calcareous slate
and limestone of the Carys Mills Formation succeeded by massive
medium grey lithic sandstone beds of the Siegas Formation. The con-
tact is an erosional one, with clear truncation of laminae in the
Carys Mills slate and about a foot of exposed releif.

Return to cars and proceed west along the highway.

43.2 Stop 5. Park along the shoulder of the highway well clear of the

road. The outcrop is in the road cut on the north side of the high-
way .

Madawaska Lake Formation. The exposure is in the core of the Pen-
nington Mountain anticlinorium. The lithology is dark grey non-
calcareous sparsely laminated slate with minor thin light grey sand-
stone beds.

211

Return to cars and turn around in order to proceed back down the
highway to the east .

45.5 Turn left on to access road to the old Highway 2. Immediately turn
left again on old Highway 2 and proceed slowly about 150 yards to a
gravelled side road. Turn right on the side road and proceed 200
yards to a large open space and park.

Stop 6. The Siegas Quarry. Park in the old processing area at the
end of the access road south of the quarry. Walk north to the quarry
and proceed to the east end of the exposure at the processing area
level .

New Sweden Formation and Siegas Formation. The eastern end of the
exposure consists of light grey, calcareous, laminated and locally
contorted slate of the New Sweden Formation. The top of the Siegas
Formation is defined at the top of the first significant sandstone
bed. The contact is conformable. The rest of the exposure consists
of almost the entire Siegas Formation in continuous sequence and
exposure for about 750 feet of section. The basal beds and the
contact with the Cary Mills Formation are obscured but outcrops of
the Carys Mills Formation may be observed in the fields immediately
to the west of the quarry. The section dips steeply to the north-
west and is slightly overturned.

There is an abudance of sedimentary features characteristic of tur-
bidite deposition from both distal to highly proximal phases. In-
cluded are graded sandstone and limestone conglomerate beds, massive
nongraded beds of sandstone and slump deposits of limestone con-
glomerate. A wide variety of cross-lamination, convolute lamination
and sole features are common.

The Siegas Quarry is the property of Atlas Construction Co. of
Fredericton, New Brunswick.

212

TRIP C-3

BIOSTRATIGRAPHIC TRIP ACROSS NORTHEASTERN MAINE

by
William H. Forbes

This trip will permit the participants to observe and
sample fossil localities from Middle Ordovician to Middle
Devonian in age across northeastern Maine. The localities
are among the best in the eugeosynclinal Appalachians and
include outcrops with shelly and/or graptolitic fauna as
well as outcrops with abundant Devonian flora.

213
TRIP C-4

LATE-GLACIAL AND HOLOCENE GEOLOGY
OF THE MIDDLE ST. JOHN RIVER VALLEY

J. Steven Kite and Harold W. Boms , Jr.
James Madison University and University of Maine at Orono

Introduction

The middle St. John River valley lies between the towns of Grand
Falls, New Brunswick and Frenchville, Maine (Figure 1). Throughout most
of this field area the St. John River serves as the U. S. -Canadian inter-
national boundary. Complete topographic coverage of the area is included
on three maps: the Canadian (1:50,000) Saint Andre and Edmundston quad-
rangles and the American (1:62,500) Stockholm quadrangle. The American
(1:62,500) Van Bur en, Grand Isle, and Frenchville quadrangles include
portions of Maine also shown on the two Canadian maps. Most of our field
work in this area was undertaken in 1977 and 1978, although this research
is currently being extended into the upper St. John River valley.

The glacial history of northern Maine and New Brunswick greatly in-
fluenced late-glacial and Holocene events. Ongoing studies on both sides
of the international border are putting together the Late Wisconsin gla-
cial history of the region (Genes and Newman, Trip B-8; Gauthier, Trip
B-9) . The chronology determined by these studies undoubtedly will add to
our understanding of postglacial events; however, in light of current
knowledge we believe two galcial events, both recognized at least three
decades ago, controlled drainage in the St. John watershed. The first
event was the formation of an ice dam in the Edmundston area which created
Lake Ennemond at altitudes of up to 680 ft in the western end of the field
area (Kiewiet de Jonge, 1951). The second important glacial event was the
construction of a large moraine across the valley at Grand Falls; after
deglaciation, this moraine dammed Lake Madawaska, which extended through-
out the middle St. John River valley at altitudes of up to 513 ft (Chal-
mers, 1886; Kiewiet de Jonge, 1951; Lee, 1963).

Lake Madawaska existed until about 10,000 B.P. when water level
rapidly declined. The emergent lake bottom was covered first by extensive
early Holocene swamps, and later by three different facies of early-to-
middle Holocene alluvium. Two of these alluvial facies have been misin-
terpreted in other studies as either glaciolacustrine or glaciof luvial
deposits even though identical sediments are accumulating in present-day
St. John River environments. The St. John River has incised its bed in
late Holocene time, leaving the old fluvial and lacustrine deposits expos-
ed in terrace scarps.

Acknowledgements

Robert Stuckenrath of the Smithsonian Institution provided 11 new
radiocarbon dates and aided in the interpretation of these dates. A

214

Grant -in -Aid of Research from Sigma Xi, and Fellowships from the Univer-
sity of Maine at Orono supported field work. Keith A. Laskowski and
Susan C. Kite provided able field assistance. This work benefited greatly
from discussions with Andrew Genes, Glenn Prescott, George Denton, Barry
Timson, Thomas Lowell, Bradford Hall, and Stephen Norton.

Stratigraphy and Geomorphology

Ice-contact stratified drift is the lowest Quaternary strata exposed
in the middle St. John River valley. This drift occurs in kames, kame
terraces, and the moraine at Grand Falls. The stratified drift is com-
posed of clasts ranging from fine silt to cobble-boulder gravel. Canadian
Shield gneissic erratics are a minor constituent. Collapse features such
as normal faults and steeply dipping beds are common. Flow till occurs
at several localities. In most exposures, ice-contact stratified drift
is capped by a layer of till; in one gravel pit the stratified drift
underlying the till layer exhibits overturned glaciotectonic folding
(Stop 10). The stratified drift presumedly accumulated during the next-
to-last glaciation of the valley, the overlying till dates to the last
glaciation. Several isolated ice-contact deposits in the field area are
not capped by till and are probably younger than the till-mantled strati-
fied deposits.

There are two distinctive tills which mantle the bedrock hills,
kames, and kame terraces in the field area. A well-compacted grey (N4:
wet color) diamicton is probably a lodgement (basal) till. This grey
till is extremely poorly sorted; clasts are predominantly limestone and
fine-grained detrital rocks with minor constituents of conglomerate,
gneiss, and volcanic rocks. Up to 90 percent of cobble and boulder clasts
are striated and there is a distinct orientation of elongate clasts. An
oxidized greyish orange (10 YR 7/4) to yellowish brown (10 YR 5/2) till,
less compact than the grey till, is widespread throughout the field area.
This oxidized till was deposited as ablation or flow till. Again, gneis-
sic erratics are minor constituents. Bedding is generally absent, but
rare layers of sand or silt occur. The oxidized till has less distinctive
fabric and fewer striated clasts than the well-compacted grey till. Des-
pite distinct differences between the two tills, we do not believe there
is sufficient evidence within the middle St. John River valley to assign
these tills to different time periods or different ice masses. Both tills
occupy the same stratigraphic position overlying the majority of ice-
contact stratified drift in the field area; hence, we believe the tills
represent different facies of the last glaciation. Older tills undoubted-
ly underlie the extensive ice-contact stratified drift, but we have not
recognized outcrops of older till in this field area.

Two outwash deposits, composed of rounded cobbles and boulders with
a sand-and-silt matrix, exist in the field area. A valley fill composed
of up to 200 ft of rounded gravel lies at the eastern end of the field
area. The head of this outwash train adjoins the Grand Falls moraine.
South of Grand Falls, the outwash occurs in a steep-scarped terrace which
is over 20 miles long. A second valley fill, 25 ft thick, is exposed in

215

I t

35 5

cO

a
o

•H

CO
C

CU
p

C
•H

Q)

42

CO

5-i
CU

>
•H

O

CO

- - ' •■■■•*•■

CD

42

H

•

cO

►»

cu

CU

p

H

CO

H

cO

t3

>

rH

CU

M

•H

CU

»P

>

•H

CU

P3

42

p

C

42

M-l

o

O

>->

P

•

CO

p

O

en

e

CU

p

H

3

X)

O

T3

42

•H

on

e^i

d

o

u

•

42

i— 1

p

cu

>•>

M

P

d

03

bO T3

•H

C

En

d

o

43

216

terraces along the Green River valley. The outwash head is a small
moraine north of the field area (Lee, 1955). Gneissic erratics occur
at both localities.

Rhythmically bedded, grey (N5) silt and clay make up a large por-
tion of the postglacial subsurface deposits in the field area. However
these lacustrine rhythmites occur on the valley floor below an altitude
of 505 ft and are not widely exposed because of overlying alluvium. Each
rhythmic bed is 0.1 to 1.0 in thick and displays graded bedding. The
lack of sand-silt rhythmites or thick basal rhythmites indicate these
beds accumulated in deep water, removed from active ice margins or deltas.
Rhythmic bedding thins and becomes less distinctive upward, uppermost beds
are massive.

We have mapped two types of lake-margin deposits in the field area:
beach ridges and deltas. Beach ridges, generally less than 5 ft high and
up to 600 ft long, are composed of well-sorted gravel and sand. All of
the known beaches are associated with Lake Madawaska strands lying between
altitudes of 463 ft and 513 ft. Unlike beaches, deltaic deposits are
associated with both the highest strand of Lake Madawaska (about 510 ft)
and with several Lake Ennemond water levels (580 ft to 680 ft). Deltaic
deposits closely resemble outwash sediments, except that deltas exhibit
steeply dipping foreset bedding, which is not found in outwash. One Lake
Madawaska delta lies at the mouth of the Green River valley outwash,
suggesting that glacial ice near the field area co-existed with the earli-
est lake strand.

Buried peat beds are exposed at two outcrops (Stop 5 and Stop 11).
Plant fragments are humified in the lower portion of both peat beds,
whereas the upper portion contains wood fragments up to 4 in in diameter.
The peat probably accumulated in a swamp that formed over Lake Madawaska
sediments after water level fell. At both localities, the peat beds
underlie alluvium of the highest alluvial terrace.

Channel-lag and channel-bar sediments are the coarsest of three dis-
tinct alluvial facies that can be identified in both modern-day sediments
and Holocene strata. These channel sediments, ranging from coarse sand
to cobble-gravel, commonly are 10 to 12 ft thick exhibiting imbrica-
tion, crude horizontal bedding, and high-angle cross-bedding. In one ex-
posure, channel gravel is indurated into a conglomerate with limonite
cement. Channel-lag and channel -bar alluvium can be distinguished from
outwash using three criteria: 1) outwash deposits must have an outwash
head; 2) coarse-grained deposits may be attributed to adjacent stream de-
position, if the present day stream is capable of transporting comparably
sized sediments; and 3) it is unlikely that the glaciers necessary to
generate outwash remained in the field area long after the climatic ame-
lioration that occurred about 10,000 B.P.

The coarse-grained channel deposits are overlain locally by grey
(N5) channel-fill silt and clay, widely exposed along the banks of the St.
John River and its tributaries. The fine-grained channel-fill lenses
range from 3 to 15 ft in thickness and some exceed one-half mile in
length. They commonly parallel the present-day stream course giving rise
to outcrops that are thousands of feet long. Bedding appears massive,

217

except in horizons rich in flat-lying plant fragments. Limonite mottl-
ing and amorphous blue masses of vivianite commonly surround plant
fragments . We have collected numerous flattened leathery mollusk shells ,
which have been leached of calcite. Present-day channel-fill basins
(such as the basin seen at Stop 9) are subaerially exposed only during
low stages of the St. John River.

Most of the alluvium cropping out in the area is floodplain sand
and silt, which may be up to 23 ft thick. Reverse grading of individual
beds is common, although sequences of beds deposited in a floodplain en-
vironment become progressively finer-grained toward the top of the se-
quence. Low-angle cross-bedding occurs in fine sand and coarse silt; angle-
of-repose cross-bedding exists in coarse sand beds. Bedding cannot be
discerned in weathered silt near the top of alluvial terraces upon which
floodplain deposition no longer occurs. Floodplain alluvium is the
uppermost (hence, youngest) strata in every exposure in the field area.

The three alluvial facies discussed above make up the highest allu-
vial terrace in the St. John River valley. This highest terrace is best
exposed in the vicinity of Siegas, New Brunswick; hence we call the land-
form the Siegas terrace. The Siegas terrace can be traced throughout
most of the reach between Grand Falls and Edmundston, at altitudes of
33 to 40 ft above low stage of the St. John River. The gradient of the
terrace surface is approximately the same as the gradient of the St. John
River prior to construction of the hydroelectric dam at Grand Falls. All
three alluvial sediments also occur in lower terraces, which lie 25 to
33 ft above low river level. These lower terraces are not paired and
probably were formed during progressive downcutting of the St. John River
channel. The present-day alluvium also includes the three alluvial facies.
Channel-lag, channel-bar, and channel-fill sediments are restricted to
low-lying areas near streams. Active floodplain accumulation, evidenced
by lack of soil profile development, occurs up to 25 ft above low stage
of the St. John.

Late-Glacial and Holocene Geologic History

The middle St. John River valley was deglaciated in stages. After
the deglaciation of the St. Lawrence lowlands at about 13,000 B.P., a
residual ice cap remained in northern Maine and New Brunswick (LaSalle,
et al., 1977). Cut off from ice domes in the Canadian Shield, this ice
cap thinned and broke into smaller glaciers. One of these glaciers stood
between Edmundston and Grand Falls while the western and eastern ends of
the field area were ice-free. This glacier probably dammed Lake Ennemond
and deposited the moraine and extensive outwash at Grand Falls. During
the highest strand (680 ft) of Lake Ennemond, the lake may have drained
through one of several potential outlets south of the field area (Figure
2A) . Subsequent lower strands (660 to 580 ft) may record the opening of
an outlet through the middle St. John River valley as the glacier wasted
in place. It is possible that an arm of Lake Ennemond extended east of
Edmundston as the glacier contracted After the glacier melted completely,

218

OTHER LATE-GLACIAL LAKES

B

MAXIMUM EXTENT OF
LAKE MADAWASKA

Figure 2. Ancient
lakes of the middle
St. John River
valley: A) maximum
extent of Lake
Ennemond and B)
maximum extent of
Lake Madawaska.
Towns are denoted
with triangles:
GF = Grand Falls,
E = Edmunds ton,
FK = Fort Kent.

219

the only obstruction of drainage in the valley was the Grand Falls
moraine and associated outwash, the drift which dammed Lake Madawaska
(Figure 2B) . In its early history the lake apparently drained through
two outlet channels at Grand Falls. One outlet paralleled the present-
day course of the St. John, whereas the other outlet was about one-half
mile south of the Grand Falls gorge (Figure 1).

The date of deglaciation and onset of lacustrine sedimentation can
be inferred from an exposure on Green River (Stop 5); here, an estimated
1200 rhythmites underlie the base of a peat bed which has yielded 5
credable radiocarbon dates of about 10,000 B.P., with an average of
9979 ± 108 B.P. If the rhythmites are assumed to be varves, and if we
assume there is no hiatus in the stratigraphic sequence, then the oldest
lake sediments accumulated at about 11,200 B.P. Deglaciation probably
occurred immediately before deposition of the lake sediments. The
highest recognized strand deposits in the vicinity of the Green River
exposure occur at elevations of about 510 ft; the altitude of these
strands suggests that the 11,200 B.P. -aged lake sediments accumulated in
Lake Madawaska .

Beaches below the highest strand record water levels during the
regression of Lake Madawaska (Figure 3) . The only well-exposed beach in
the field area (Stop 10) overlies a shallow-water silt deposited while
other beaches formed at higher altitudes. As the water level fell, the
emergent lake bottom was covered by swamps. The 10,000 B.P. -aged peat
bed at the Green River exposure (altitude: 474 ft) formed after Lake
Madawaska dropped about 36 ft below its highest nearby strand. This 36 ft
regression occurred in about 1200 years, indicating an average water level
drop of 3 ft/ 100 years. A 9,900 B.P. -aged peat bed near Parent, Maine
(altitude: 446 ft; Stop 11) suggests that the lake level fell 28 ft with-
in the century after 10,000 B.P. Indeed, the radiometric control over
the Parent and Green River exposures does not preclude the possibility
that the 28 ft drop in lake level occurred catastrophically. Our field
data suggest the decline after 10,000 B.P. was so rapid that no mappable
beach deposits formed below the 10,000 B.P. strand. The southern outlet
to Lake Madawaska was probably abandoned during the rapid drop in lake
level. Organic-rich channel-fill sediments taken from the abandoned out-
let yielded a rabiocarbon date of 9830 ± 160 B.P. (GSC-56).

We have considered a number of possible causes for the rapid demise
of Lake Madawaska. Three of the more probable causes include isostatic
uplift in the upper St. John watershed, failure of a drift dam or ice
dam, and changes in the hydrology of the watershed induced by climatic
amelioration. Research in the upper St. John valley hints that other ex-
tinct lakes in the watershed had similar histories to Lake Madawaska.

Wood incorporated into channel gravel indicates that a river flowed
over the emerged lake bottom by 9820 ± 130 B.P. (GSC-18), or possibly
somewhat later if the wood was reworked prior to deposition. Lake
Madawaska ceased to exist before this event, although small discontinuous
lakes may have persisted for a short time in lower areas of the valley.
The early Holocene St. John River deposited gravel at altitudes as low

220

S

co"

/ w

/ °J
/ °>

s

1 ^™*

♦1

/o

o

/ <-

00

o

1 z

I

8

• © © •

UJ

<

\-

<

I

<

UJ

<

o» ©

©

\

CO
UJ

d uj

00
CO <

2j

UJ
■Htt

o> „

Cft ©

z
<

a:
*

X
O

CO

u.

o

UJ

-J

u_
o
a:
q.

■ ►o

8

ujujujuj

O (BCD CD

LO
CO
LO

^

E

o

O

LO O

o

LO

£ LO

LO

3annriv

LO

CO

I

E

lO
CVJ

UJ

CO

co

O Q

<
a:
o

o
u-

UJ

o o
CO

Q

O
lO

■8

o
o
<3-

T3

CO cu
CU CO

d id hi

•H ,0 ,C

3

4-1

co

CO
•H

CO

01

C

•H

fa *0

• C

PQ cd

U

O 4-)

O CO

<y»

o ,fi

ana

4-1

T3 O

C X

CVJ ^-^

cd

2
<

• O
fa

1 I

CO H
•H
O 4-1
O

o •*

* CO
O 4->
O

cd

CD
H

O
O

cu

c

•H

u
■u
co

a
o

c

•H

60
C
•H
fa

U

i

cu

4-1

o

U-l
rH

cd

I

cu

c
o

CO

3

60

c

•H
fa

u

§

o

•H

4-1

cd

4-1

CO

o

CO

•H

cd

XI

4-1

CO

cu
d

•H
rH

c

cd cd

V4 r-i

4J -■*.

CO 4-1

cd

cd cu

M fa
co

cd u-i

£ o
cd

T3 CU

eg (jo

S cd

oj g

5 o

cd M

. CU

ro V4

M

CU CU 4J ^

M <4-l -H

60'H (4 0
•H cd CN

fa CU

a)

a

•H

O T3
•H
4-1

CI

cd
a. u
e u
3 co

CO

CO •

cd fa

>, PQ
U
cd ^— *

u
cd

a

o

221

as 427 ft. The newly formed reach of the river had an extremely low
gradient over the former lake bottom. Immediately after the draining of
the lake the St. John River and its tributaries began an aggradational
episode, which was widespread by 8200 B.P. (Figure 4). Aggradation con-
tinued until after 4250 ± 50 B.P. (SI-3703) . By the end of the aggrada-
tional episode a broad floodplain existed with a surface that sloped
from 480 ft near Grand Isle to 458 ft near Grand Falls. Small, coarse-
grained alluvial fans, a few feet above the floodplain surface, formed
near the mouths of many tributaries.

Downcutting of the St. John River began once the river attained a
graded profile. The floodplain and alluvial fans were incised to form
the Siegas terrace. Lower terrace-like surfaces fall into two age
groups. Weathered surfaces lying 5 to 15 ft below the Siegas terrace
are alluvial terraces abandoned during the late Holocene. Unweathered
surfaces 10 to 30 ft below the Siegas terrace are part of the present-
day floodplain.

Geoarcheology

We did not find any artifacts in our field work, but the strati-
graphy of the middle St. John River valley suggests potential for
significant archeological sites. Lake Madawaska persisted long after
the arrival of man in Maine and the Maritimes, as evidenced by the Paleo-
Indian site at Debert , Nova Scotia, which is dated 10,589 ± 47 B.P.
(Stuckenrath, oral communication, 1979). Some of the Lake Madawaska
beaches apparently are contemporaneous with the Debert site, and signifi-
cant Paleo-Indian sites may be located on, or just above, these relict
lake shores. The alluvial stratigraphy of the Siegas terrace spans most
of the period between 10,000 B.P. and 5000 B.P., during which time there
is little known about man in Maine and the Maritimes. The Siegas terrace
probably contains sites that could fill this void in the archeological
record. Sites of this age within the Siegas terrace may have a well-
developed stratigraphy in which each artifact assemblage is temporally
separated from other assemblages. Late Holocene artifacts probably occur
in surface sites on the Siegas terrace, surface and buried sites on young-
er alluvial terraces, and buried sites on the present-day floodplain.

References

Chalmers, R. C., 1886, Preliminary report on the surface geology of

New Brunswick: Geology and Natural History Survey of Canada Annual
Report 1885, part GG, 58 p.

Kiewiet de Jonge, E. J. C. , 1951, Glacial water levels in the St. John
River valley (Ph.D. dissertation): Worcester, Mass., Clark
University, 116 p.

222

m

00/

► O ♦ □

S

1
u

,— *

CO

♦

2

o

ro

Q.

_l_

oo

/ *0

V)

o

/ °J

tL

k>

1 0>

^~

1 ™"»

CO

+i

in

o

(VI
OD

la
1 **

+i

0)

m

I

|2

s

►

/Q-"

D

1

/s

\0z

o

\z

j I

ro
1

0

/ ~3

(0

*-*

1

/»-'

Q.
00

ro
m
ro

/ °°

o
m

3C

/ **.

' o

1

UJ

lu
o
o

Ld
_l
Q
Q

o
m

(M

0.
03

O
O
CJ
•M
O

m
c\»
oo

Ld
-J

U_
O

Q:

oo

oj

C\J i

<

Q
00

P Q

<

8

E
-o-

<t IT)

Ld

C E

o o

o m

m —

LU
O <_>

'I

00
Q

O

m

o

ID

in o

& m

aaniinv

m

E
in

CJ

o
o

*T3

PL.

OJ

•

C

PP

•H

CO

m

4-)

00

,a

o

+ 1

OJ

o

u

v£>

<u

i—i

5

v£>

05

*■

0)

/>— s

4J

1— 1

CO

O
CO

5-i

1

0)

rH

A

CO

■u

">w^

o

<U

p-l

OJ

•

r-l

PQ

43

H

o

•

+

(3

O

m

•H

o^

•P

r^

CO

m

TJ

CO

U

• •

•

00

E

•-N

60

3

<

CO

•H

i-H

>

o

(U

3

On

C

rH

CO

0)

rH

1

a

CO

H

o

CO

rH

C

^^

o

•H

33

CO

•

rH

PM

a)

a

•

rH

T3

pq

T3

o

13

o

o

•H

H

o

s

<4-l

^H

o

43

+

■u

O

•H

o

>>

S-<

r^-

iH

1

r~-

)-i

O

co

cO

•H

W

a

T3

CO

G

00

CO

■r

^

<f

O

#>

a)

QJ

O

n

43

o

a

•U

CT>

GO

en

•H

6

i

PJ4

o

i— i

u

CO

m

^— '

223

LaSalle, P., Martineau, G., and Chauvin, L., 1977, Morphology, strati-
graphy, and deglaciation In Beauce-Notre-Dame Mountains-Laurentide
Park area: Ministere des Richesses Naturelles (Quebec), DPV-516,
74 p.

Lee, H. A., 1955, Surficial geology of Edmundston, Madawaska and

Temiscouata Counties, New Brunswick and Quebec: Geological Survey
of Canada Paper 55-15, 14 p.

1963, Field trip guide for the Friends of the Pleistocene 26th

annual reunion, Riviere-du-Loup , May 25-26: Quebec, 29 p.

Mileage

Itinerary

0.0 Assemble at entrance to Centennial (Municipal) Park, Grand

Falls, New Brunswick, Saint -Andre 1:50,000 quadrangle. Start-
ing time 8:30 A.M. Eastern (9:30 A.M. Atlantic). Walk west
300 yd to bridge over St. John River.

0.0 Stop 1. Walk to center of bridge, facing waterfall.

Grand Falls at the confluence of St. John and Little Rivers.
The St. John drops over 120 ft within one mile of the hydro-
electric dam; most of the drop occurs at this waterfall. The
view is most spectacular during April or May, when spring melt-
water increases discharge to as much as 1.65 million U.S.
gallons per second. Carefully walk to the other side of the
bridge to view gorge. The erosive power of the river during
peak discharge is evidenced by the large potholes and the tree-
less zone that extends far up the sides of the gorge. Before
10,000 B.P., the path of the gorge was the northernmost of two
Lake Madawaska outlets. The southern outlet, 900 yd from this
bridge, was underlain by over 200 ft of drift in the Grand
Falls moraine (Figure 1). The drift-floored southern outlet
was abandoned about 10,000 B.P. The northern outlet probably
was drift floored at 10,000 B.P., also; if it had been bedrock
floored, the St. John River would have become entrenched in
the more easily eroded southern outlet. Hence, the deep gorge
beneath the bridge was apparently carved after 10,000 B.P. be-
cause of superposition of Lake Madawaska' s northern outlet.

Return to assembly point, proceed in vehicles northwest on
Victoria St.

0.2 Turn left on Broadway.

0.6 Turn left on Condon St. The railroad tracks visible south of
the intersection overlie the abandoned southern outlet of
Lake Madawaska.

224

0.8 Cross railroad tracks.

0.95 Turn left on River Road at 3-pronged intersection; cross
railroad tracks again.

2.35 Turn right at entrance to large gravel pit.

2.35 Stop 2. Park well out of truck traffic.

Grand Falls outwash. This coarse gravel was deposited by
meltwater from the glacier that built the Grand Falls moraine.
Cobble and boulder lithologies include Canadian Shield gneisses
which originated west of the St. Lawrence, it is possible that
these erratics were not transported across the St. Lawrence by
the Grand Falls glacier but were reworked from older drift.
Large trough cross-bedding was exposed in the southern face of
the pit in 1977 and 1978. A veneer of sand and silt caps the
coarse gravel. These finer sediments probably mark a reduction
in meltwater after the glacier retreated from Grand Falls.

Return to vehicles, turn left on River Road.

3.75 Turn right on Condon St., immediately after crossing railroad
tracks .

5.2 Turn left onto Route 2 (Trans-Canada Highway) west.

9.4 Bridge over Canadian Pacific Railroad tracks; note the view to
southwest across St. John River. Landforms include modern-day
floodplain, the Siegas alluvial terrace (site of several farms
and paved road) and a till-mantled bedrock ridge. These land-
forms are shown on both Saint-Andre 1:50,000 and Van Buren
1:62,500 quadrangles.

11.6 Note the well -developed segment of the Siegas terrace on this
side of the St. John River.

16.5 Intersection of Route 2 and Route 17, continue west on Route 2.

18.4 Cross Grande Riviere (Grand River).

18.5 Stop 3. Park in front of gate on left side of road. Follow
trail to river bank 300 yd south of gate.

Three alluvial facies of a late Holocene terrace. At least
6.0 ft of gravel crop out at the base of this exposure. The
gravel unit displays sedimentary structures (crude bedding and
imbrication) and grain-size distribution similar to present-day
channel-lag and channel-bar deposits; the gravel in the exposure
probably accumulated in a channel facies. 4.6 ft of massive
silt and clay overlie the gravel. Sedimentary characteristics
include abundant carbonized plant fragments concentrated near
bottom of unit and limonite mottling. The depositional environ-
ment was probably similar to the modern-day channel-fill facies.

225

The channel-fill clay and silt are covered by 3.6 ft of sand
and silt which were deposited in the floodplain facies.

Return to cars. Proceed west on Route 2.

21.0 Cross Riviere Siegas (Siegas River).

21.2 Note the unpaved road to left which leads to the abandoned site
of Siegas Station. The Siegas terrace, the highest alluvial
terrace in the field area, is best developed in this vicinity.
An excellent exposure of the three alluvial facies of the Siegas
terrace occurs on the Canadian bank of the St. John River, 400
yd southwest of Siegas Station. This exposure can be seen from
Stop 11 and from U.S. Route 1 in Maine. The deposits are
similar to those exposed at Stop 4. We will not have time to
visit this exposure; continue west on Route 2.

24.7 Enter Edmundston 1:50,000 quadrangle in the town of Ste-Anne-
de-Madawaska .

25.2 Stop 4. Turn right at last intersection before bridge over
Riviere Quisibis (Quisibis River). Follow right turn with a
quick left. Park car on dead end road paralleling Route 2.
Walk across bridge and proceed to exposure on bank of Riviere
Quisibis 125 yd southwest of bridge.

Alluvial facies of the late to middle Holocene Siegas terrace.
Again channel gravel, channel-fill silt and clay, and floodplain
sand and silt are exposed. The contact between the channel-fill
and floodplain facies is marked by alternating beds of sand and
fine silt. Anorganic-rich zone at the base of the channel-fill
alluvium yielded wood dated 8250 + 200 B.P. (W-353, Lee, 1955).
During low-water levels, silts, possibly deposited in Lake
Madawaska, are exposed below the channel-gravel unit. Another
exposure of the same units occurs on the bank of the St. John
River 400 yd southeast of this locality.

Return to vehicles, continue west on Route 2.

25.3 Cross Riviere Quisibis.

31.0 Cross Green River (Riviere Verte).

31.2 Turn right on gravel road located a few yards east of exit to
town of Riviere Verte (Green River Station). Continue north-
east between brick house and campsites; proceed as far as road
conditions permit.

32.0 Stop 5. Park before reaching shallow gravel pits; walk to first
gravel pit.

Holocene point-bar and floodplain deposits. This operation is
extracting gravel that is nearly identical to point -bar

226

sediments accreting along Green River (Stop 6). Much of the
gravel appears to be derived from an outwash plain and a delta
(altitude: approx. 510-515 ft) formed at the mouth of the out-
wash plain north of this locality.

Walk southeast under power line. Cross Green River in vicinity
of mid-channel bar. (During high water, fording the stream can
be avoided by approaching this outcrop from the east along the
power line right of way. However, this eastern approach re-
quires a one-half mile foot traverse through poorly-drained
terrain. )

Green River exposure (Figure 5) . This is the best exposure of
Lake Madawaska sediments in the middle St. John River valley.
The compact grey diamicton at the bottom of the exposure dis-
plays a distinct orientation of elongate clasts: long axes
trend N 60° W. Canadian Shield gneisses are a minor constituent
of this diamicton which is probably a basal till. The compact
grey till is overlain by a highly deformed unit that includes
gravel, sand, silt, and clay. Portions of this unit contain
rhythmic beds indicating glaciolacustrine deposition. Glacio-
lacustrine deposition of the deformed unit is also supported
by the fact that this unit underlies a thick sequence of rhy-
thmically-bedded clays and silts which were deposited in Lake
Madawaska. There are approximately 1200 rhythmic beds in this
exposure. The upper one-half foot of lake sediments does not
display rhythmic bedding, possibly because they were deposited
in shallow water. Assuming the rhythmic beds are varves (i.e.
deposited annually) Lake Madawaska sedimentation lasted for
about 1200 years at this locality. The oldest beds of this
lacustrine unit are probably the same age as the highest lacus-
trine beaches and deltas in this vicinity (altitude: approx.
510 ft). The base of the peat overlying the lake sediments has
yielded five credable *^C dates with a mean age of 9979 ±
108 B.P. and a median age of 10,090 ± 130 B.P. If there was
no hiatus between lacustrine and peat deposition the
oldest lake sediments date to about 11,200 B.P. Hence, this
locality was probably deglaciated shortly before 11,200 B.P.
The top of the peat bed dates to 9285 ± 70 B.P. (SI-3706); after
this date the peat was buried under the fining upward alluvial
sequence exposed at the top of the outcrop.

Return to vehicles, turn around and drive toward Route 2.

32.7 Stop 6. Park in front of campsites, before reaching Route 2.
Walk 250 yards southeast to west bank of Green River.

Present-day point -bar and floodplain deposits. The west bank
of Green River consists of a point bar. Note the coarse grain
size and clast imbrication. Total thickness of the gravel,
including gravel below river level, is about 20 ft.

227

cr

ill

2

g

3

Q.

O

OD

o

EL

0-0.

0,' 0. 0.

CD

CD CD

Cb CD CD

8

OO

eae

(O

—

♦ 1*1

+1 -H *l

O

O*

Cvj (0 JO

3

&|

5> el*

i

m

r, <T>" #
O ff> *

<->

i
o

hi S> 3S

o

19

(oSo)

UJ

o

4-1

a

cO

4-1
CO

cu

B
rfl
CO

S
o

V4

<4-l

13

cu
C

•H
CO

4J

£>
O

CO

Q)
4J
CtJ
T3

d

O

u

CO
O
O
•H
T3
CO
5-i

X
•H
CO

"4-4

o
cu

4J

CO
13

CO

d
o

rH
CO
B

o

C
CO

CU

XI

4J

TT3
CU

T3

.H
CU

•H

CO

CU

CO
CU

«4-l Pl,
O

CU

>

•H

CU

S-i

CO
O

X!
cu

5-1

CU

b

Pi
d

CU

cu
o

m

PQ

CO

4-1

•H
CO
O

a
cu

T3

5-i
CU

-a

B
o

u

n3
CU
>

•H

5-1
CU

CO
•H
M

cu

4J

CO
6

o

•H

d

cO
00

5-1

o

M3
CU

M
U

$

cu

5-i

o
o
o

O CU

-h -a

d

d

3
O

u

CO

5-i
CU
4-1
CO

3

rH
O

CJ

CO

O

5-i

a

13 0\

cu on

X> CO
CO

4-J I

CO M

cu zr>

a cu •

04t fH

m

m
cu

5-1

d

PQ

O
m

<4-l

o

B

4J

O O

•H ^H

Pm CU -

X. CN

228

Cross bridge to east bank of Green River, north of Route 2.
Examine the sand and silt of the floodplain facies. Nearly
every bed displays reverse-grading : within each bed the grain
size is silt at the base and gradually increases to sand at the
top. The contact between individual beds is abrupt. The
mechanism producing reverse-graded bedding is unclear.

Return to vehicles, continue toward Route 2.

32.8 Turn right (west) on Route 2.

37.5 Turn left on Route 14.

41.0 Cross Riviere Iroquois (Iroquois River).

41.5 Stop 7. Park on right side of Route 14, north of radio tower.

Lake Madawaska beach. A low ridge in the lawn north of Route 2
is one of the highest beach deposits known between Edmundston
and Grand Falls. The beach has an altitude of 512 ft and was
probably formed at the same time as the delta along Riviere
Verte, early in the history of Lake Madawaska. A cross section
of a beach ridge is exposed at Stop 10.

Continue west on Route 14.

42.6 Route 14 passes over an abandoned channel of the Madawaska
River.

43.3 Cross Madawaska River; look for signs giving directions to
Madawaska, Maine or to Customs.

43.5 Following directional signs bear right, then bear left.

43.6 Turn left on Rue de St. Francois.

43.7 Turn right onto International Bridge; enter Frenchville 1:62,500
quadrangle (this leg is also shown on Edmundston 1:50,000 quad-
rangle) .

44.0 Stop at U.S. Customs, Madawaska office.

44.1 Turn right on U.S. 1.

48.3 Cross Gagnon Brook.

48.6 Stop 8. Turn left into Madawaska Sand and Gravel pit, park well
out of truck traffic.

Gagnon Brook delta. This compound delta was built into ice-
dammed Lake Ennemond. Several water levels ranging from 680 ft
to 580 ft are recorded here, the best developed delta topsets
occur at about 620 ft. Relict ice-wedge casts may be exposed in

229

these topset beds. The Ruisseau Felix -Martin (Felix Martin
Brook) compound delta can be seen in New Brunswick, one mile
southwest of this locality.

Return to Route 1, turn right.

58.6 Enter Grand Isle 1:62,500 quadrangle (also shown on Edmundston
1:50,000 quadrangle).

59.3 Turn left on unpaved road.

59.6 Stop 9 . Park on top of terrace;- do not attempt to drive down
terrace scarp. Walk to first culvert.

Cyr's Ponds: abandoned channel sediments. The silt and clay
accumulating in these water bodies are analogous to the massive
silt and clay of the channel-fill deposits. Comparison of
grain-size distribution, sedimentary structures (especially
massive bedding and limonite mottling), and faunal assemblage
shows the sediments accumulating in these basins are nearly
identical to those exposed at Stops 3, 4 and 11. The abandoned
channels are not true ponds but are backwater basins in which
water level is controlled by the height of the St. John River.
The channel-fill sediments are exposed only during low water
levels. Floodplain sands and silts accumulate on higher sur-
faces nearby.

Return to vehicles, turn around and drive to Route 1.

59.9 Turn left on Route 1.

70.9 Enter Stockholm 1:62,500 quadrangle (this area is not shown on
any 1:50,000 Canadian quadrangle).

72.8 Note unpaved road on right, immediately before passing over
knoll east of Parent, Maine. Continue southeast on Route 1.

72.9 Turn right on next unpaved road, just east of knoll crest.

73.0 Stop 10. Park near entrance to gravel pit. Walk to southwest
end of pit .

Till-mantled kame underlying beach built during regression of
Lake Madawaska. The southwest face exposes ice-contact strati-
fied drift which underlies an oxidized buff till. Glaciotec-
tonic isoclinal folding was exposed in 1977 near top of ice-
contact stratified drift. One to two feet of shallow water
lacustrine silt covers the till. At the southern corner of the
pit the silt underlies a beach ridge composed of subrounded
gravel and sand. The beach crest has been stripped away. The
existance of higher beaches nearby, coupled with the presence
of shallow-water silt under the beach gravel demonstrate that
this beach formed during regression of Lake Madawaska.

230

Return to Route 1 .

73.1 Turn right on Route 1.

73.8 Stop 11. Park on right side of road across from white trans-
former. Walk 100 yd northeast to St. John River bank.

Parent exposure (Figure 6) . At least 700 rhythmic beds of Lake
Madawaska silt and clay are exposed near river level; the basal
contact of this lacustrine unit is not exposed. Radiocarbon
dates obtained from samples taken along contact between lacus-
trine sediments and overlying peat bed indicate that the level
of Lake Madawaska dropped below this elevation (446 ft) about
9900 B.P. Peat deposition ended about 8855 ± 65 B.P. (SI-3704).
The massive grey silt and clay overlying the peat bed is similar
to sediments previously identified as channel-fill deposits.
However, these massive sediments are not underlain by channel
gravels. These sediments probably accumulated in a flood basin
that formed when the St. John River began building levees during
the early Holocene aggradation. The uppermost horizons in the
exposure are floodplain deposits. The youngest radiocarbon date
obtained from organic lenses in the floodplain sediments is
4250 ± 50 B.P. (SI-3703). Floodplain deposition on the upper
surface of the Siegas terrace ended after this date.

Return to vehicles. End of field trip. Those returning to
Presque Isle should continue southeast on Route 1, taking care
not to miss the right turn in Van Buren.

231

Ld

Z>
10

O

Q-
X

u

o

£

o

o

cr
u.

Q

Ld
N

_J
<

or

Ld

Z
Ld

o

Ld

Z>
(/)

O

CL
X

LU

I-
Z
UJ
01

I- 'X(-, hii

ao

C\J

cc
>

O

CO

UJ

tfi

CO

ro

5

UJ
Q

CO

o

&
w

g

o

CO

0)

•H
Pn

232

TRIP C-5

SEDIMENTOLOGY OF SILURIAN FLYSCH, ASHLAND SYNCLINORIUM, MAINE

by

David C. Roy-
Department of Geology and Geophysics
Boston College

The Silurian of the Ashland Synclinorium offers an opportunity to study
turbidity current deposition in a sequence that may well represent continental
margin sedimentation following the Taconian Orogeny (Roy, 1978). The uplift of
pre-Silurian rocks to the west and northwest of the synclinorium produced emer-
gent terrain with a narrow shelf environment and apparently steep eastward
slopes into a basin in which sedimentation was uninterrupted. A 2600 meter
(8500 feet) section of Silurian shale, graywacke, and conglomerate was deposit-
ed at the base of the slope and in the basin. Basinward from the former land
area the deposits become generally finer-grained and more distal in turbidite
characteristics. Upward in the Silurian section, similar changes reflect both
reduction of land relief and progressive advance of marine conditions onto the
detrital source area during the Silurian. These changes produced increasingly
more distal conditions everywhere within the basin.

The stratigraphy of the Stockholm region is summarized in Figure 1 and is
discussed by Roy and Mencher, (1976) and Roy, Trip B-6, this volume). On this
trip we will focus on turbidite deposition within the Jeratland Formation (STOPS
2,4, and 6) which forms part of the upper half of the Silurian section. The
lower and more proximal part of the section as seen in the Frenchville and
New Sweden Formations will be seen at STOPS 3 and 7; the pre-Silurian strati-
graphic setting (Madawaska Lake and Cary's Mills formations) will be examined
in STOPS 1,3, and 5. The locations of all stops are given in this volume in
Figure 2 of Roy, (Trip B-6).

Post-Taconian Submarine Erosion

We will visit an exposure of the Taconian angular unconformity between
the Ordovician Madawaska Lake Formation and the Lower Silurian Frenchville
Formation (STOP 3) . The details of the exposure are given in the stop descrip-
tion below and the regional aspects of the Ordovician-Silurian transition are
given elsewhere in this volume by Roy, (Trip B-6). It is concluded that at
least the final pre-Frenchville erosion at this exposure was submarine and it
is possible that all of the folding and erosion of the Madawaska Lake beds were
accomplished below sea level,

Submarine unconformities with apparantly substantial hiatuses in eugeo-
synclinal regions are not widely reported as such. Direct and indirect evidence
of intraformational submarine erosion with short hiatuses are of course widely
observed and reported in both shallow and deep-water sedimentary sections. In
the present case the unconformity is angular and is a systemic interface that
represents a non-trivial hiatus. The erosion is considered to have occurred
on a post-Taconian submarine slope for four reasons:

1) To the west and northwest, Lower Silurian rocks are absent,

233

AGE

COLUMN
(Thickness in Meters)

LITHOLOGY AND FEATURES

ENVIRONMENT

NOMENCLATURE

Green and red silty shale and slate
, with interlayered thin beds of very
calcareous, laminated fine-grained
graywacke and siltstone.

Thinly interlayered gray shale, non-
laminated silty shale, laminated
silty shale and fine-grained cal-
760 careous graywacke. Medium-grained

and coarse-grained graywacke common.
Turbidite structures typical of
graywacke beds. Abundant graptolites,

0 )

> 0

V

+»

■J

V *

•C P <

3 «

3 Tj

tj

u

r ♦*

■* -• J

h a t>

§

* y

t*

H n X

S # a

0

0

> v

c

« «T

JS J3 «]

OJ

to

!>. >

V "

«U T) %

0.

CM

c

«t

/ l-l

JO C .

> U t)

0

-d

*• y

•J

O CO S

•OB

0

c

no <

* V

« a 0 •

O

0

«

v *>

> *" >

3 -H P «

I—

0

/ 01

>> M j

> O B W J4

O

■p

<h y

0

m 0 f

(1 CO

0

n V

N

UK C

t. • O a)

•H

■p ~

^ -P

to > X

« « >. >

1-*

r •"•

>> >

> O ♦> T* >,

rH y

CD

•a m X

H «) 1 ■)

0

3

c *« X.

«j h a t<

ft.

V <*

11 to \

U a Z «0

+1

o
o
eo

Unconformity
(Locally Angular)

Green-gray slate
with lesser
graywacke

Gray slate thinly
interlayered with
micritic limestone
and minor graywacke
("ribbon rock") ;
graywacke more
abundant in lower
part of section.

Distal

Flysch

Intermedi

at

e

and

Distal

Flysch

•h

CO

0

§ »

■H C

«M T3

01

^ 03

O C

c to

I, ^

CO

0 a

O

n

•h eg

O. «»

C 0) 0)

4J <M

c

O C -P

u

tl -H

i-* ai 1-1

0 V

+> »"

•P «M 01

Q. C

a) «J

I. 0

^-*

^ B

ova.

•H hi

-d .0

0. c oj

0 «J

V =>

■w TJ

0 B

B 01

.-( I.

•rt ,0

In

«J CO c

X 3

ti «-

P S -H

0 co

+J O

en .0 en

L.

c

•H 3 00

ft.

M

O 01 J3

Submarine \

Erosion \

Distal

Calcareous

Distal

Flysch

Flysch

Fogelin
Hill
Format ion

Jemt land
Format ion

French-
Iville Fm

Figure 1:

Stratigraphic section of the Stockholm- Jemtland area
(from Roy and Mencher, 1976).

whereas above the unconformity and to the east a thick "deep-
water" section representing almost the entire Silurian is
present.

2) Paleocurrent measurements and provenance of sandstone beds in
the Frenchville and Jemtland formations indicate downslope
transport of detritus from the west and northwest.

3) Frenchville sandstone and conglomerate beds contain brachio-
pod assemblage composed of representatives of mixed depth
communities (Roy, 1973; McKerrow and Ziegler, 1971).

4) At least the final erosion of the Madawaska Lake Formation
produced large flute molds on the erosional surface which
suggests differential scour of semiconsolidated mud by
sediment-laden currents.

234

EXPLANATION
GRAYWACKE BEDS
In+ervol C :

Complex Ripples

r*""*^ Simple Ripples

Convolu+e

Lamination

Interval Bs

Simple Lamination

Politic Lamination

*' Interval A:

□

PELITIC BEDS

Laminated
Siltij Shale

Si 1+4 Shale
Slate

t Direction of
Grading

|^>.| Pelite Fragments

N

Paleocurrent

px Measurement

SJ Flame Structure
r>^ Covered

DCROY

Figure 2; Lithologic and depositional description of the Jemtland
section at STOP 4. This is section 2 in Figures 4 and 5.

235

Early Silurian Sedimentation

Following the Taconian uplift to the west, clastic debris was shed east-
ward into a broad basin. Along the western margin of the basin, at the position
of the Ashland Synclinorium, a thick sequence of sandstone and conglomerate was
deposited (Frenchville Formation) which gave way to more shale-and limestone-
rich facies out in the basin (New Sweden and Spragueville formations) . These
Lower Silurian litho facies are more easily examined in the Ashland-Presque Isle
area and are described in more detail elsewhere in this quidebook (Roy, Trip B-6)

Where we are able to see the base of the Frenchville along the western
margin the Ashland Synclinorium the formation rests unconformably on the
Madawaska Lake Formation as just described in connection with the exposure at
STOP 3. The more eastern lithofacies each rest conformably on the "ribbon rock"
limestone of the Upper Member of the Carys Mills Formation. It is therefore
inferred that the Taconian hiatus decrease eastward across the Ashland Syn-
clinorium to essentially zero and that the Taconian uplift can be fairly well
dated by the influx of mud that terminated the more cyclic "ribbon rock"
deposition of the Carys Mills. The change from Cacys Mills sedimentation to
those of the New Sweden and Spragueville formations is pretty well dated as
about graptolite zone 19 of the Early Silurian (Pavlides, 1968; Roy and
Mencher, 1976). It is likely that the Aroostook-Matapedia Basin in which the
Carys Mills was deposited became more shallow as a result of the Taconian up-
lift that raised land to the west (Roy, 1970a). Subsidence of the basin (and
possibly also the western upland) during post-Taconian sedimetation is inferred
from the great thickness of the Silurian section in the basin and the fining
upward of the section.

Jemtland Sedimentation

The Jemtland Formation is a thin-bedded flysch composed of five principal
rock- types and two minor rock- types. The major rock-types are:

1) Indistinctly laminated calcareous , platy, silty shale. This
rock type contains the most abundant and best preserved grap-
tolites.

2) Nonlaminated, calcareous, silty shale. This rock type usually
shows an irregular bedding cleavage.

3) Dark-gray to black fine-grained thinly cleaved slate. Cleavage
in this rock type is usually axial planar to the folds, or
nearly so.

4) Light-gray, calcareous, micaceous, laminated and cross-
laminated, qiiartzose, siltstone or fine-grained sandstone
(graywacke) .

5) Gray, medium- to coarse-grained, calcareous, feldspathic and
lithic graywacke.

236

These rock types are interlayered on the scale of centimeters and tens of centi-
meters. It is common for lithic graywacke to grade upward into the micaceous
quartzose graywacke in thicker graded beds. The minor rock types are tuff
(STOPS 2 and 6) and micritic limestone.

All Jemtland exposures can be described litho logically in terms of the
above rocks (usually just the major rock- types) and can be described sediment-
ologically in terms of a Bouma-type turbidite model. Roy (1970a, 1970b) has
shown that the pelitic rock types (types 1,2, and 3 above) do not show any
preferred ordering with respect to each other or the graywacke beds and hence
must be treated separately in terms of sedimentation. The graywacke beds may
be nicely described in terms of Bouma's (1962) A,B, and C intervals. A pelitic
bed relates only indirectly to the turbidity current that deposited the under-
lying graywacke bed as discussed below.

Figure 2 shows a graphical description of the outcrop at STOP 4 (essent-
ially the "type" Jemtland) in terms of Bouma's model as modified by Walker
(1965, 1967). This exposure and that at STOP 2 are typical of the graywacke
facies of the formation. Figure 3 shows two sections (similarly described)
from the more eastern slate facies of the Jemtland; see Figure 9 for the loca-
tions of these two eastern sections. From all three of the sections the
following generalizations are evident:

1. Amalgamated (composite) turbidites (Walker, 1965) are common
in the graywacke facies.

2. Basal "A" intervals are much more common in the graywacke facies.
North of the latitude of Washburn there is in fact a fairly
regular eastward decrease in the proportion of graywacke beds
that begin with basal "A" intervals as shown in Figure 4.

3. The transition from "A" to overlying "B" intervals is usually
gradational whereas transitions from "B" intervals to "C"
intervals almost always involve an erosional interface. The
upward ordering of A,B, and C is statistically favored (Roy,
1970b).

4. Grading is common in "A" intervals and is present but less
obvious in many "B" intervals. Graded "A" intervals were
produced by deposition under conditions of very rapid fallout.

Non-graded "A" intervals may be material transported by grain
flows .

5. Both of the silty shale rock types are more abundant in the
graywacke facies than the finer-grained dark slate. The dark
slate increases in abundance basinward as shown in Figure 5.

6. The most common type of pelite to follow the graywacke beds is
the type that is the most abundant in the particular section
(Roy, 1970b).

In addition there are features that are not well displayed in the
graphical sections of Figures 2 and 3 but are generally observed in outcrop:

2t H

(Colcite
25

23

+

e 'vem

I

103

m

<n

83

SI

A

r£

>r

%
«

1

5*>»J II

81

si

H

! 8<<

83

80

73

W>^»-

:

101

WJ

100

I;

is

B

25

lit

115

111

DC ROY

237

Figure 3: Lithologic and depostional descriptions of outcrops of the

slate facies of the Jemtland Formation: A) Outcrop at
section 4 of Figures 4 and 5; B) Outcrop at section 5 of Figures 4 and 5
(see also Figure 9) .

238

7. Graywacke beds overlying either of the two types of silty
shale described above generally have erosional sole features
(flutes, tool marks, ridges and grooves, etc.) whereas soles
of graywacke beds overlying the fine-grained pelite show
load features (e.g. load casts, flame structures, etc.).

8. The clay/silt matrix of the graywacke beds appears to be pri-
mary. This conclusion is based on depositional segregation of
the matrix in sedimentary structures and the sharp, little-
altered boundaries of feldspar and rock fragments in many

of the graywacke beds. Carbonate alteration is more common
than clay alteration in the graywacke.

9. Bioturbation is very rare and delicate siltstone laminae within
pelite beds are traceable laterally for great distances as is
well shown at STOP 2.

10. Laminae within MB" intervals become thinner, more traceable,
and more clay/ silt rich upward. Pelite laminae are commonly
seen in the upper parts of "B" intervals. Weisbrich (1977)
and Robert Brown (unpublished data) have produced parallel
laminated beds in flume experiments at plane-bed flow velocities
in which laminae thickness and, to some extent, clay/ silt
incorperation are found to depend on the rate of bed aggrada-
tion. The experiments were run by allowing sand to fall-out

at known rates through a recirculating slurry of fixed lutum
(clay plus silt) content. As a rule rapid bed aggradation
(high fall-out rate) produces indistinct and thick laminae
whereas low bed aggradation (low fall-out rate) produces thin,
well defined, and more laterally persistent laminae. Incorpor-
ation of lutum in the experimental deposits increases with
decreasing bed aggradation rate with marked increases to a few
percent at aggradation rates less than about .05 cm/min. So
far, pure pelite laminae within experimentally produced lami*
nated fine sand beds have not been formed and their presence
in "B" intervals of Jemtland and other turbidites is still a
mystery. Variations in laminae thickness and persistence
within "B" intervals, however, appears to be useful in quali-
tatively assessing variations in bed aggradation rates during
deposition of these turbidite intervals. The experimental
aggradation rates used by Weisbrick and Brown to produced
lamination comparable to that seen in the Jemtland ranged from
.003 to .5 cm/min; thus the deposition time of "B" intervals
of the Jemtland were possibly tens of minutes in duration.

11. Ripple lamination in "C" intervals typically represents climb-
ing ripples with low angles of climb and thus indicate low
values of the ratio of bed aggradation-to-ripple migration
(Jopling and Walker, 1968; Scheible, 1979). This ratio is
given by: Ar/Mr = tan Ac where Ar is the bed aggradation rate,
Mr is the ripple migration rate and Ac is the angle of climb
(Scheible, 1979). In fairly well developed climbing ripples,
angles of climb of 5 - 10° (Ar/Mr of . 1 to .2) appear necess-
ary to permit clear discernment of climbing ripple cosets;

239

NumWar, spr««4>and
a**coa« erf pakocurntnt
mMNUttwttnts ;arT oufaad
traUcatas cunvm Sanaa

© (SI. 3)

Section studied Per-cant of grauuxxclai
m detail bads u>it+> A-mrtrvala

Figure 4: Variation in proportion of graywacke beds in the
Jemtland Formation that have basal A- intervals.

/I T 2 *°

i ffjjy I

240

Number, spread,and
averaqeof paleocurrent
measurements ;arrou)head
indicates current sense

©

Section studied
in detail

(R.8)

Percent of pel itic beds
that is slate

DCB0Y

Figure 5j Variation in proportion of fine-grained slate in the
Jemtland Formation.

241

somewhat higher angles of climb are probably necessary in more
complicated natural cosets to clearly see the climbing charac-
ter of the ripples. The generally low Ar/Mr ratios in
Jemtland "C" intervals probably is the cause of the erosional
surface at the base of most such intervals. Non-erosive trans-
itions from "B" to "C" intervals would require high Ar/Mr
ratios or nearly vertical angles of climb to prevent substrate
erosion. Scheible's (1979) work suggests that as a practical
matter realization of such high Ar/Mr ratios means that flow
velocities may have to be below velocities for initiation of
motion of the sand grains on a flat bed (to greatly reduce Mr)
while appreciable sand of that grainsize is still available in
the flow to fall-out. It has been shown that ripples are stable
and can migrate at flow velocities less than that required to
initiate ripples on a flat bed (Middleton and Southard, 1977).
Such low flow velocities, however, may inhibit development of
ripple bed forms from the precurser flat bed conditions at the
close of "B" interval deposition. This is because the initia-
tion of ripples on a flat-bed requires flow velocities on the
order of the flat-bed initial motion velocity unless there is
significant perturbation of the flat-bed (Middleton and South-
ard, 1977). It is also likely that the presence of fall-out
from a turbidity current reduces the lower velocity limit of
upper flat-bed stability to values well into the ripple
stability field as established (for a given grainsize) under
conventional non-aggradation experimental conditions. During
turbidity current deposition, ripple development is probably
suppressed both by fall-out and the continuing reduction of
grainsize during sedimentation (Allen, 1970).

12. Pelite clasts are common in the coarser-grained graywacke beds.
These, usually tabular, clasts may have maximum dimensions
that are on the order of the thickness of the graywacke bed
itself. Most of the pelite clasts are fragments of Jemtland
shale added to the detritus by submarine erosion upslope from
the position of deposition. It is common that the clasts are
positioned within the graywacke bed at levels of 1/4 or 3/4
of the bed-thickness above the base of the bed. This same
localization of pelite clasts is found in Frenchville turbi-
dites and the present writer has no very good explantion for
this tendency.

The above features suggest, or are at least consistent with, deposition
of the graywacke beds from lutum-rich turbidity currents. The clay and silt of
each turbidity current cloud probably remained in suspension for some time
following sand deposition. Upward mixing of the warm shelf-water of the turbi-
dity current after reaching the cold bottom water of the basin probably helped
to maintain a fairly persistent near-bottom suspension cloud which was re-
supplied by subsequent turbidity currents. The unlaminated silty shale and the
pelite of the fine=grained slate probably represent deposition from such more
or less stagnant suspension clouds. The non-laminated silty shale would
represent deposition from the near-source part of the cloud.

242

/

cu

i

X ~^~9k%
AN Sfv^€&*

/ A

/

N

' ;-*/-,,

0m\

/

/

/

/

/

/

/

Sfc

/

/

~20OO'

;1 / // sj >oci/DCR

Figure 6:

Geologic map of the Blackstone Siding area
showing the locations of STOPS 2 and 3. Base
from an uncorrected airphoto.

The laminated silty shale is more difficult to explain because the fabric
of the shale and the common orientation of graptolites within it suggest
current effects during deposition. It is possible that the laminated silty
shale beds were deposited from the silt-rich phase of basinal suspension clouds
that were moving in response to contour currents, residual turbidity currents,
or other currents within the basin.

Post-Jemtland Sedimentation and the Close
of the Silurian

Along the axis of the Stockholm Mountain Syncline (Roy, Trip B-6; Figure
2) the Jemtland Formation is overlain by the Fogelin Hill Formation which may
span the Siluro-Devonian boundary (Roy and Mencher, 1976; Figure 1). The
Fogelin Hill is conspicuous in its content of red and green slate and absence
of the coarser-grain graywacke beds found in the Jemtland. The red and green
slates suggest important changes in the geochemical environment of the basin
following Jemtland deposition. The basin appears to have become more oxidiz-
ing and more distal with respect to detrital sources.

243

The Fogelin Hill Formation appears to span the time interval of the
Salinic Disturbance for which there is good evidence both to the east (Presque
Isle area) and to the west (Fish River Lake area) . Rocks very similar to the
Fogelin Hill are present in the Portage area to the west of the synclinorium
where they contain Early Devonian graptolites and rest unconformably on
Ordovician rocks (see Figure 2 of Roy, Trip B-6; Roy and Mencher, unpublished
data). Crinoidal limestone is present as thin beds in both the Portage area
and high in the Fogelin Hill section within the synclinorium. The Fogelin Hill
Formation is interpreted by Roy, (this volume) to have been deposited in a
residual basin of the synclinorium that survived but was restricted by the
Salinic "uplifts" on both sides. .

References

Allen, J.R.L., 1970, The sequence of sedimentary structures in turbidites,

with special reference to dunes; Scottish Jour. Geology, v. 6, pp. 146-161

Bouma, A.H., 1962, Sedimentology of some flysch deposits; Amsterdam, New York,
Elsevier, 168p.

Jopling, A.V. and Walker, R.G., 1968, Morphology and origin of ripple-drift
cross-lamination, with examples from the Pleistocene of Massachusetts;
Jour. Sed. Pet., v. 38, pp. 971-984.

Middleton, G.V. and Southard, J.B. , 1977, Mechanics of sediment movement;
Soc. Econ. Paleon. Min. Short Course No. 3.

Pavlides, Louis, 1962, Geology and mangamess deposits of the maple and Hovey
mountains area Aroostook County, Maine; U.S. Geol. Survey Prof. Paper
362, 114p.

1968, Stratigraphic and facies relationships of the Carys Mills Formation,

northeast Maine and adjoining New Brunswick; U.S. Geol. Survey Bull. 1264,
44p.

, Neuman, R.B. and Berry, W.B.N. , 1961, Age of the "ribbon rock" of

Aroostook County, Maine; Art. 30 in U.S. Geol. Survey Prof. Paper 424B,
pp. B65-B67.

Rast, N., and Stringer, P., 1974, Recent advances and the interpretation of

geological structure of New Brunswick; Geoscience Canada, v.l, pp. 15-25.

Roy, D.C., 1970a, The Silurian of northeastern Aroostook County, Maine (Ph.D.
thesis); Cambridge, Massachusetts Institute of Technology, 483p.

1970b, Variations in a thin-bedded turbidite sequence from the Silurian
of northeastern Maine; Geol. Soc. Amer. Abs. with Programs (An. Mtg.),
v.2, no. 7, pp. 669-670.

1973, The provenance and tectonic setting of the Frenchville Formation,
northeastern Maine; Geol. Soc. Amer. Abs. with Programs, v. 5, p. 214.

244

1978, Tectonic History of northeastern Maine; Geol. Soc. Amer. Abs. with

Programs (N.E. Section Mtg.), v. 10, no. 2, p. 83.

and Mencher, E. , 1976, Ordovician and Silurian stratigraphy of northeast-
ern Aroostook County, Maine, in Page, L.R., ed., Contributions to the
stratigraphy of New England; Geol. Soc. Amer. Mem. 148, pp. 25-52.

Scheible, M.A. , 1979, An experimental study of climbing ripple lamination and
its use in determining depositional conditions (M.S. thesis): Chestnut
Hill, Boston College, 105p.

Walker, R.G., 1965, The origin and significance of the internal sedimentary

structures of turbidites; Yorkshire Geol. Soc. Proceedings, v. 35, pp. 1-32.

1967, Turbidite sedimentary structures and their relationship to proximal

and distal depositional environments; Jour. Sed. Petrology, v. 37, pp. 25-43.

Weisbrich, G.J., Experimental studies of Bouma B-zone lamination (M.S. thesis):
Chestnut Hill, Boston College, 65p.

Itinerary

Mileage

Assembly point is the parking lot of Keddy's Motor Inn. Starting time
is 8:00 A.M. Head north on U.S.I through Presque Isle. Mileage starts
at the bridge over the Aroostook River north of town.

0 Bridge over Aroostook River. U.S.I from here to Caribou passes through
some of the best potato country in Aroostook County. The bedrock
supporting the agriculture is composed of the Carys Mills and Sprague-
ville formations.

7.75 Bear left off "new U.S.I" onto "old U.S.I" toward downtown Caribou.

11.05 Turn left onto Route 161 in Caribou just beyond the downtown shopping
mall.

11.35 Turn left toward Main Street.

11.40 Turn right on Main Street which is also Route 161.

12.25 Bear right at junction with Route 228 and continue north on Route 161.

16.55 STOP 1. Park on shoulder of highway. Please be careful since the
exposure is on the inside of a curve.

245

This is one of the best exposures of the Lower Member of the Carys
Mills as I have mapped it in the Caribou Quadrangle. Graptolites
from this exposure have been assigned by Pavlides and others (1961)
to zone 13 of the Ordovician (Caradocian) . The strata here are about
3500 feet (1200 meters) below the base of the Upper Member of the
Carys Mills which is composed of the famed limestone "ribbon rock"
(largely of earliest Silurian age) of the Aroostook-Matapedia Belt.
Four main rock-types are present here: 1) gray, fine-grained, green-
weathering, calcareous slate; 2) light-gray, medium- to- fine grained,
micaceous, calcareous, graded, laminated/ cross laminated quartzo-
feldspathic graywacke; 3) light-gray, massive, micaceous, rusty-orange
weathering, calcite-veined quartz graywacke in which sedimentary
structures are rare; 4) dark-gray, pale-gray weathering, laminated
micritic limestone. This pre-Taconian sequence is inferred to have
been deposited in an oceanic basin receiving abundant terrigenous
detritus from intraoceanic sources which may include volcanic islands
to the west and a tectonic island arc in the present Miramichi
Anticlinorium (Rast and Stringer, 1975). Terrigenous sources are
inferred to wane in importance during deposition of the Upper Member
of the Carys Mills were calcareous turbidites dominate the deposit.

Return to cars and continue north on Route 161.

19.15 Bear left at Y- intersection. Outcrops of the Late Early Silurian New
Sweden Formation are on the left side of the highway. The "Caribou
Station" of Boston College's seismic network is located a short dis-
tance straight ahead from this intersection (so much for geophysics
on this trip) .

20.00 Turn left (west) at the light in Sweden, Maine (Texaco Station).

22.65 Continue straight at the sharp right (north) turn in the paved road.
We are entering a major active lumbering area and you are urged to
watch out for large trucks; it is further recommeded that you let
them have the right-of-way.

24.15 Turn right (north). Road becomes a bit narrower and rougher here
(usually) .

25.65 Turn left (south) onto the entrance road for the "Blackstone Siding
quarry".

25.85 Park as best you can.

NOTE: We will spend about 1.5 hours in the Blackstone Siding area
(Figure 6) examining STOPS 2 and 3. The outcrop at STOP 3 is small
(!) and is best examined by small groups so it is suggested that
cars go to it one or two at a time (following directions below)
during our stay here. Those more interested in studying Jemtland
turbidites may simply stay at STOP 2.

STOP 2. The Blackstone Siding quarry provides the best exposure of
the Jemtland Formation in the region. It displays the most proximal
phase of the formation and provides slightly weathered pavement sur-
faces that permit careful study of turbidite features. The strata

246

strike about N35°E, are essentially vertical, and top to the southeast.
No reversals of facing have been found. In the southeast corner of
the pit are exposures of basal beds of the Aquagene Tuff Member of
the formation which can be traced discontinuously to the northeast
for a distance of about nine miles (14 km). Graptolites are ubiqui-
tous in this quarry and some fine specimens are possible. The strata
here are of Early Ludlovian age (graptolite zones 33-34 of the
Silurian) .

Return to cars and head out to the main road.

26.05 Main lumber road. Turn left (west).

26.15 Blackstone Siding (International Paper). In this vicinity we cross

the Jemtland-Frenchville contact that is not observed here but is seen
to be gradational in river exposures along strike to the northeast.

26.30 Little Madawaska River.

26.3 Road to right (north); stay on main road.

26.8 Turn right onto a segment of the old haul road that is now been
bypassed.

27.3 STOP 3. In this small outcrop (some have called it an "incrop") the
Taconian unconformity is exposed between the Madawaska Lake Formation
(Ordovician) and the Frenchville Formation (Figure 6) . The contact is
interesting because it is angular, there are large flute casts on the
sole of the basal sandstone bed of the Frenchville, and step-lineations
on the sole of the sandstone preserve the intersection of Madawaska
Lake bedding and the erosion surface. The relationships of these
features are shown in Figure 7A. The flute casts are assymmetric with
the steep slope of the flute located on the down-dip (pre-Frenchville)
side of the flute and the "floor" of the flute mold was essentially
parallel to the bedding in the underlying Madawaska Lake Formation
(Figure 7B) . Taken together, these features suggest Taconian folding
(submarine?) of semiconsolidated Madawaska Lake beds followed by
submarine erosion by a turbidity current and finally deposition of the
pebbly lithic graywacke. The material of the graywacke bed may have
been transported by grainflow or other mechanisms, arriving at the
bottom position (canyon floor?) represented by the outcrop after the
turbidity current had largely passed.

Return to cars and continue northwest along the old stretch of the
lumber road.

27.40 Main road. Make sharp left to head back (eastward) on the main road.

27.5 As you go up the hill the Madawaska Lake-Frenchville contact is again
crossed (we are only a short distance along strike from STOP 3) .

28.30 Little Madawaska River.

28.45 Blackstone Siding.

247

- a)

-</)

<u
en 4J

P-. 00

o c

H -H

co £

O

4J 42

cd co

a) co

00 u

a co

3 CO
t-h a
a
a)

•P *J

co 3
cd rH
O M-l

I

(1) T3

p a
3 cd

rH

<+4 00

a

cd 'O •

cu a)

C cd

O CU U

•H ^ 4J

■u cd co

id iJ J3

cd co

co cu

cd j:

u

<u

I

cu cd
co cd

o

(3
•H

00

c

T3
CU

id

CO T3

C cu

O 43
•H

4J O

Cd -M

CO C!
CX CU
•H

n

o

43
CO

a
o

CU

cd

•H
H >
<U 4^

U
U

CO

co
cd

o co

C cd

4J cu o

c n

<U tn cu

CD I 43
CU CU +J

P-l pL, LM
O
I I

<! w ft

•H

43
co

o

•H
4-1

cd

cu
u

a

00
•H
fa

cu
u

248

28.55 Entrance to Blackstone Siding Quarry. Continue east.

30.05 Turn left (east).

31.55 Continue straight on paved road.

34.20 Intersection with Route 161 (Texico Station). Turn left (north).

38.80 STOP 4. Park as directed. The exposure is on the west side of Route
161 in a slight curve. Care must be exercised in both parking and
crossing the highway. This will probably be a good place for a LUNCH
STOP. A small store (if open on Sunday) in Jemtland (near the stop)
may provide provisions.

This is the "type exposure" of the Jemtland Formation. The formation
here is similar to that seen at STOP 2 but is more slate-rich (less
silty shale) and is inferred to be slightly more distal. The outcrop
is near the middle of the Jemtland and is on the southeastern flank of
the Stockholm Mountain Syncline (see Figure 2 of Roy, Trip B-6, this
volume). The proportion of graywacke (58 percent) is similar to that
in the sequence at STOP 2. Figure 2 shows the details of the strati-
graphic section here. Common sole features readily permit paleocurrent
measurements which show a west-to-east flow of the turbidity currents
(Figures 2,4, and 5).

Continue north on Route 161.

39.50 Approximately here we cross the axis of the Stockholm Mountain Syncline ,
Youngest unit is the poorly exposed Fogelin Hill Formation.

40.75 Little Madawaska River. Frenchville Formation is exposed in the river
upstream from the bridge.

43.30 STOP 5. Park on the shoulder of the highway.

This is one of the better exposures of the Madawaska Lake Formation
(Ordovician) in the "type area" of the formation. In this exposure
the rocks are typical and consist of predominant olive-green, fracture
cleaved slate with lesser thin beds and laminae of calcareous quartzo-
feldspathic siltstone or fine-grained sandstone. Here also are thin
beds of rusty-weathering pyritiferous micritic limestone. One exposure
here also shows red slate which is a minor but conspicuous rock type
in the formation near the unconformity with the overlying Frenchville
Formation in this area. The Madawaska Lake is the temporal equivalent
of the Winterville Formation (volcanic rich) to the west. More
westerly in the Madawaska Lake, lithic graywacke increases in abundance
presumably reflecting proximity to volcanic islands.

Continue north on Route 161.

43.45 Turn around in the Rest Area and head south.

46.15 Little Madawaska River.

249

46.20 Turn left (east) on paved road.

47.70 Turn left (north) at Stockholm Post Office. Continue north across the
Little Madawaska River and R.R. tracks, by the store, and up the hill.

48.60 Entrance to Stockholm Town Dump on left. Enter and drive to refuse

area and park. Town permission is required to enter this scenic area.
Walk southwest along the strike of the lithic tuff horizon that is
more or less exposed in the dump.

STOP 6. Aquagene Tuff Member (provisional name) of the Jemtland
Formation. An almost complete section of the tuff sequence is avail-
able in these brush covered exposures and is depicted in Figure 8.
The member (here 63 feet thick) is stratigraphically in the middle of
the formation on the northwest flank of the Stockholm Mountain Syncline
as shown in Figure 2 of Roy, (Trip B-6, this volume). The member is
thinner and more discontinuous on the southeast flank of the fold.
Fossil localities stratigraphically below, above, and within the member
in the immediate vicinity date the tuff as Ludlovian of the Late
Silurian. Four broad tuffaceous rock types comprise this section.
Type I tuff, the most abundant, is light gray-green, chalk-white
weathering, and variably lithic. In section devitrified shards and
pumice fragments are clearly visible. Plagioclase, quartz, and chlo-
rite are dominant mineral phases. Lithic tuff, designated as Type II,
is composed of sand-to pebble-size fragments of intermediate and mafic
volcanic rocks, plagioclase, quartz, myrmekitic quartz-feldspar, and
detrital carbonate (including fossil fragments) set in a matrix of
devitrified shards and pumice fragments. Type II tuff usually transi-
tions upward into Type I in apparently graded sedimentation units.
Type III tuff is a chlorite-rich rock that differs from Type I pri-
marily in the presence of large chlorite "patches" that are flattened
parallel to a bedding parting. Upward gradation from Type I to Type
III tuff is observed in three beds (Figure 8) . Type IV tuff is fine-
grained, light gray-green in color, cherty in appearance, and occurs
in thin beds less than 6 inches thick that commonly show faint
lamination. This type forms a 7-foot interval in the middle of the
section. In thin section tuff of Type IV shows devitrified shards and
pumice fragments in a dominant matrix of microcrystalline quartz-
chlorite-plagioclase. The sequence of tuff here is thought to have
been deposited in deep water and probably consists largely of resedi-
mented tuffaceous material. The ash was most likely derived from the
west where contemporaneous volcanic rocks interlayered with shallow-
water sediments are present.

Return to cars and return to main road.

48.9 Main road at entrance to the dump. Turn right.

49.8 Intersection at the Stockholm Post Office. Turn right (west).

51.3 Intersection with Route 161. Turn left (south).

53.2 Outcrop of STOP 4 on the right.

57.8 Intersection at light (Texaco Station). Continue south.

250

".-. -.-..■* -M

Cm.>.Jt.^w,j.>.-l

Quartz. Vein

^^

* • * . * ■',••••*■■" •"•.*!

r/vt/tfWV

4hMn

i •.• • • • • %* At*

■■iiii.i'.i.'i.i'ir.'ii'.'i

^;t,;;,;/:'"i-yM^

■1

•'■'•V'

?©' .

EXPLANATION

V.!.' It

IX

TYPE I TUFF
TYPE H TUFF
TYPETJI TUFF

TYPE H TUFF

THINLY BEDDED
SLATE, SHALE
AND GRAYWNCKE

COVERED

0 1 2 3 4 5 b 7

■ * -* * * i i '

FEET
VERTICAL SCALE

Figure 8: Stratigraptiic section of the Aquagene Tuff Member of
the Jemtland Formation at STOP 6. The types of tuff
are described in the stop description.

251

60.4 Turn right (southwest) onto gravel road.

61.05 Turn right (west) onto paved road.
62.25 Railroad crossing in Colby.

63.3 Turn a hard left (southeast) onto gravel road at "T" intersection.

64.4 Turn right (west) onto paved road (Route 228).
65.35 Turn left (south) onto gravel road.

67.75 Turn right (west) onto paved road.

69.15 Junction with Route 228 in Perham. Turn left (south).

69.35 Main intersection in Perham. Turn left (southeast) and park in
church parking lot (if vacant) .

STOP 7. In the village of Perham there are at least three mappable
horizons of manganiferous ironstone within the New Sweden Formation.
These horizons and others in the Perham area are shown in Figure 9.
Here the ironstone is interlayered with gray, calcareous, finely
cleaved, laminated, phyllitic slate with lesser micritic limestone
beds. Facing here is to the southeast. The ironstone is thinly
laminated and generally quite calcareous. The details of laminae
mineralogy have not been determined as yet, but hematite-rich (includ-
ing specularite) , carbonate-rich, and manganese-rich varieties seem
most prominant. The manganese- rich laminae are reported to contain
predominantly Braunite C3Mn 0 . MnSiOj by Pavlides (1962);
Braunite is also found in small ovoid masses as well as in laminae.
The origin of the laminated ironstones is not well understood. They
appear to be primary sedimentary deposits which have been altered by
diagenesis and low-grade metamorphism. They occur at various strati-
graphic levels within the New Sweden and are rarely more than a few
feet thick and a few hundred feet in strike-dimension. They are
commonly interrupted by "barren" slate intervals. The New Sweden is
a deep-water formation, approximately 4000 feet thick, that is distal
to the turbidite-rich phases of the Frenchville Formation. The iron
and manganese have a logical source in the volcanic rocks of the
Ordovician Winterville Formation which formed the large emergent
terrain (Taconia) a "short" distance to the west. I rather suspect
that the laminated ironstones are explicable in terms of turbidity
current transport of shallow, warm, shelf water (plus sediment) , with
the iron and manganese in both particulate and soluble forms, into a
deep-water basin. The details await a better understanding of the
laminae mineralogy and trace-element chemistry.

The road log ends with STOP 7. Participants may return to Presque
Isle by continuing south on Route 228 to Washburn and taking Route 164
to Presque Isle. Route 164 intersects U.S.I just north of Presque Isle
Presque Isle is about 18 miles from here. Those wishing to go to
Caribou simply return to the intersection at mileage 69.15 and turn
right (east); you will pass through Carson and join Route 164 in Jacobs
and thence to Caribou. Caribou is about 12 miles from here.

252

14-1

o

PQ

5

o

T3

43

C

CO

cd

CO

<

CU

£3

CO

•H

C

iH

o

•H

T3

•u

CU

O

u

CU

4-1

CO

o

13

'O

c

• M

cd

CO

a

r^

o

N

PU

•H

O

M

H

O

CO

43

«4-l

CO

O

3

o

CO

M

c

cu

o

M-l

•H

•H

4-1

a

•

Cd

cd

O

O

60

4-1

o

c

o

H

cd

43

6

Q*

<D

U

43

M-l

•H

■M

O

cd

60

CO

T3

c

CU

CU

•H

o

4->

5

cd

O

O

u

CU

42

4-1

S-i

CO

u

cu

o

Cd

4-1

o

cu

cd

£3

L4

o

3

cd

•H

T3

G

6

a

cd

CT3

•H

43

CO

V-i

CO

•H

cu

■V

PL,

X

CU
CO

QJ

m

cd

xl

o

m

4J

CO

(4-4

cu

•

O

c

CO

•H

cd

a

hJ

cu

cd

LJ

e

cd

o

CO

T3

•H

CU

60

cu

u

O

>-i

cd

i-H

3

cu

O

60 rH

cu

•H

o

o

to

M-l
O

• •

c^

CO

CU

cu

3

Lj

•H

3

rH

60

4-1

•H

3

to

O

253

TRIP C-6

WISCONSINAN GLACIATION OF EASTERN AROOSTOOK COUNTY, MAINE
Andrew N. Genes, William A. Newman, and Thomas B. Brewer *
Boston State College, Northeastern University, Boston State College, Boston

Introduction

The purpose of this trip is to examine the Van Buren and Mars Hill
Drifts which constitute the surface glacial deposits of eastern Aroostook
County and relate them to models of late Wisconsinan glaciation. In addition,
moraines extending from St. Francis through Caribou to Fort Fairfield, and
moraine complexes at Mars Hill and Lower Macwahoc \j±11 be examined ( Fig. 1 )
We will use the field evidence that relates to the problem of the origin of
the drifts and the moraines to suggest deglaciation models.
Maps: Mattawamkeag, Wytopitlock, Sherman, Mars Hill, Caribou, Presque Isle,

15 minute series.

Late Wisconsinan Glaciation

The late Wisconsinan ice advance had reached its maximum extent onto
the continental shelf and had begun to retreat by 17,000 years BP. (Connally
and Sirkin, 1973 ). Between 13,500 and 12,500 years BP. the active ice
margin had retreated from the present coastal position leaving a belt of
submarine moraines ( Borns , 1966, 1967, 1973; Stuiver and Boms, 1967;
Stuiver and Borns, unpub . data ). The general recession that produced the
coastal moraine complex was interrupted by a major readvance in eastern
Maine that culminated in the sea at Pineo Ridge approximately 12,700 years
ago ( Borns, 1967 ) .

While the receding ice margin was depositing submarine moraines along
the present coast of Maine, the Champlain Sea transgressed up the St.
Lawrence Lowland from the Gulf of St. Lawrence reaching the vicinity of
Ottawa, Ontario, by 12,800 years BP. ( Richard, 1978 ). This Champlain
Sea penetration cleaved the Laurentide Ice Sheet along the trend of the
St. Lawrence Lowland. The cleaved ice sheet to the south of the Champlain
Sea formed a residual ice cap over southern Quebec, Maine, and New Bruns-

*Work supported by the Maine Geological Survey

254

ES KERS

STRIAE

\

y younger
o older

MORAINE
) observed
inferred

1 St. FRANCIS

2 WINTERVILLE

3 POR TAGE

4 CARIBOU

5 Ft. FAIRFIELD

6 STOCKHOLM

7 CASTLE HILL

8 SOUA PAN

9 MARS HILL

10 LOWER MACWAHOC

11 ORIENT

i\ Riviere Verte
/pt^ -^FrencMlle '_<^*G_rond Isle

.' .' -Von Burin
^ \ \^ V '• .' ,' • '•' •••'.•. ij^Stoeknolml • Gro

rond
ills

Ashland*

\Va

Castle Hill

Mars
Hill

ORIFT

VAN BUREN
MARS HI LL

A

«

« V

* v

km

50

K

\

Fig. 1. Generalized surficial geology of Aroostook County

255

wick after 12,800 years BP. ( Boms and Hughes, 1977 ). This represents
the minimum date at which the Laurentide ice in Maine became detached from
the Laurentide ice north of the St. Lawrence River.

Deglaciation of the newly separated ice cap proceeded through down-
wasting and recession. Residual ice cap conditions in northern Maine, and
probably adjacent New Brunswick, lead to the formation of occassional
moraine complexes, outwash deposits, and eskers . Peat overlying till
along the Green River in New Brunswick suggests that northern Maine was
deglaciated sometime before 10,200 years BP. ( Kite, 1979).

Glacial Drifts

Two distinct late Wisconsinan glacial drifts mantle eastern Aroostook
County ( Genesand Newman, 1978, 1979 ). The Van Buren Drift, covering the
very northern portion of the county ( Fig. 1 ), includes a compact, silty,
buff to dark brown till containing 3-5% clasts of Precambrian granite gneiss.
The Mars Hill Drift covering southern Aroostook County and beyond, includes
a loose, sandy, brown till. Clasts in this drift include DeBoullie Grano-
diorite, Chapman Sandstone, and Mapleton Sandstone - lithologies character-
istic of regions immediately to the northwest. Granite gneiss inclusions
have yet to be observed in this till by the authors.

Directional indicators point to emplacement of both tills by ice
moving in a southeasterly direction but no areal or stratigraphic contact
between these tills have been found.

Moraines

The moraines of Aroostook County are generally low, narrow, undulating
structures that are indestinguishable at times from bedrock controlled
topography and irregularities of dense forest canopy. The Lower Macwahoc
Moraine was discovered only after a clear cutting operation permitted a
view of the surface morphology. The moraine complex at Caribou closely
approximates the southern margin of the Van Buren Drift and consists of
Van Buren Till.

The orientation and position of the moraines, meltwater channels,
outwash, and eskers indicate that glacial recession was accomplished by

256

downwasting and recession from central Maine to northern Aroostook County.

Glaciation Models

In light of the available data, two distinct models for the late
Wisconsinan history of northern Maine are proposed to explain the observed
relationship between the Van Buren and Mars Hill Drifts :

a) Separate Ice_Cap Model: A separate ice cap developed in central
Maine during mid Wisconsinan time which eventually expanded and
interacted with the advancing Laurent ide Ice Sheet from north

of the present St. Lawrence Valley. Eventually they coalesced and
the Maine ice cap was displaced to the southeast. The Van Buren
Till was lodged beneath the Laurentide ice and the Mars Hill Till
was lodged beneath the Maine ice.

b) Frozen_Bed -_Melted Bed Model: The late Wisconsinan Laurentide ice
cap extended across Maine and onto the continental shelf. Both the
Van Buren and Mars Hill Tills were lodged by this ice advance. The
presence of tills containing Canadian Shield erratics (Van Buren)
or local rock types (Mars Hill) was a function of the thermal
regime that existed at the ice-bedrock interface.

References

Boms, H.W., Jr., 1966, An end moraine complex in southeastern Maine
(abs.): Geol. Soc. America Abs. for 1966, Spec. Paper 101, p. 485.

, 1967, Guidebook, Friends of the Pleistocene, 30th Ann,

Reunion, Machias, Maine: Orono, Maine, Maine Univ. Press, 19 p.

, 1973, Late Wisconsinan fluctuations of the Laurentide

Ice Sheet in southern and eastern New England: Geol. Soc. America
Mem. 136, p. 37-45.

Borns, H.w. , Jr., and Hughes, T.J., 1977, The implications of the Pineo

257

Ridge readvance in Maine: Geogr. phys . Quat . , v. XXXI, no. 3-4,
p. 203-206.

Connally, G.G., and Sirkin, L.A., 1973, Wisconsinan history of the Hudson-
Champlain Lobe: Geol. Soc. America Mem. 136, p52-53.

Genes, A.N., and Newman, W.A. , 1978, Two late Wisconsinan ice frontal pos-
itions, northeast Maine: Geol. Soc. America, Abs . for 1978.

Genes, A.N., and Newman, W.A., 1979, Late Wisconsinan glaciation of
Aroostook County, Maine: Geol. Soc. America, Abs. for 1979.

Kite, J.S., 1979, Postglacial geologic history of the middle St. John River
Valley: Unpub. M.S. Thesis, Univ. of Maine, Orono.

Richard, B.H., 1978, Age of Champlain Sea and "Lampsilis Lake" episode in
the Ottawa-St. Lawrence; in Current Research, Part C, Geol. Surv.
Canada, Paper 78-1C, p. 23-28.

Stuiver, M. , and Boms, H.W., Jr., 1967, Deglaciation and early postglacial
submergence in Maine (abs): Geol. Soc. America, Abs. for 1967, Spec.
Paper 115, p. 59-60.

Itinerary

Mileage

0 Assembly point is parking lot at Keddys Motor Inn, Presque Isle.
Starting time 8:00 A.M. Go north on route 1. Follow signs to
Caribou.

9.6 Take left on 161 north and continue on route 161.

18.9 Take right off 161 and head up hill.

19.6 §£2E_1.l Caribou Moraine. Here, gravel pit operations have exposed
a section 10-12 meters high permitting us to examine the morph-
ology and geographic relationship of the moraine to bedrock hills
Note the approximately 3-5% occurrence of granite gneiss clasts

258

in the exposure. These erratics are derived from the Laurentide
region of Quebec. South of this locality we have yet to find
granite gneiss in tills. This moraine, is part of a discontinuous
moraine complex which extends from St. Francis through Caribou
and Fort Fairfield, and into New Brunswick where it is correlated
with the Grand Falls Moraine.

Return to cars promptly and return the same way as entered.

20.2 At foot of hill take left onto route 161 south.

22.9 Turn left into Caribou Country Club off 161 south.

Stop_2^ Caribou Country Club. Drive to the top of the ridge by
the club. From this vantage point we have a good view of the
morphology of the moraine. Low undulating ridges extend northerly
toward the cut seen on stop 1. Although the ridge has been smoothed
it is a natural feature in excess of 12 meters thickness . Cuts
behind the locality expose granite gneiss inclusions within
the till comprising the moraine.

Return to cars.

23.4 Return to 161 south and turn left. The bedrock ridges are per-
pendicular to the drift ridges in this region. From this point
southward no granite gneiss inclusions have been found in Aroostook
County tills.

26.3 Take right on route 161. Head toward Presque Isle.
27.0 Take left onto route 1.

27.2 Take right to Presque Isle on route 1. We will follow the beautiful
Aroostook River Valley. Once we leave Caribou and head south toward
Presque Isle the topography is smooth and regular reflecting
bedrock control. Note the clumps of boulders in the fields. These
are purposely left by the farmers so that they can identify or
locate areas where bedrock is close to the surface and thus avoid
damage to their plows.

38.3 Cross the Aroostook River.
39.8 Northland Hotel on left.

40.6 University of Maine at Presque Isle on right.

48.7 Turn left off route 1 south toward Westfield. We are observing a
change from bedrock controlled topography to hummocky constructional
topography. This is the beginning of the Mars Hill Moraine ,

50.3 Note the cut into the pit on the right as we swing left.

259

50.5 Turn right. We have entered the Mars Hill Moraine complex.

50.8 Note Carry's Mills bedrock exposure on right. This is one of the
few bedrock exposures in the area. Mars Hill is seen directly
ahead.

51.9 Turn right at the "T". All topography in this region is con-
structional. As we continue south note the increased frequency
of moraines being cut by rills and meltwater stream channels. On
the right is Green Mountain Ridge. Mars Hill and Green Mountain
Ridge funneled late Wisconsinan ice thus constricting its flow
resulting in the formation of the Mars Hill Moraine complex.

55.3 To the right is a former glacial meltwater channel.

55.7 Take left on secondary road by large partially dissected moraine
ridge.

55.8 Cross route 1A onto road directly across highway.

55.9 Stop 3. Mars Hill Moraine Complex. Park cars along the road.
The thickness of the moraine ridges varies from 0-30 meters. The
complex is an irregular, hummocky sheet consisting of outwash,
kames , kettles, and smaller till hummocks. All of these forms
have been dissected by outwash streams that flowed directly from
a receding ice margin. Clasts in this drift include the DeBoullie
Granodiorite, Chapman Sandstone, and Mapleton Sandstone-all local
lithologies. Clasts of Mars Hill Conglomerate within the till appear
only at the southern end of the till complex and indicates lodgement
by till with a southern flow. An extensive moraine complex char-
acterized by distant moraine ridges rises above the general level

of the drift surface. At Mars Hill, Squa Pan Lake, Macwahoc, and
Orient these ridges take the form of frontal moraines, some of
which are associated with outwash. The Mars Hill Moraine complex
is the most striking. It extends from Mars Hill(495m elevation)
westward to Green Mountain Ridge (360m elevation) and thence
northward approximately 16-24 kilometers to Easton and Phair.

Return to cars and go straight ahead.

56.5 Take right fork. Cemetary on left. We crossed a large former melt-
water channel just prior to the road intersection.

57.3 T-junction. Take right fork. Note possible lateral moraine
flanking Mars Hill.

60.2 Left off main road.

60.6 Take left onto L.L. Boyd and Son's Farm.
Stop 4. Boyd Pit.

260

This exposure in the Mars Hill Moraine is near the distal terminus
of the moraine. To the south of this ridge the region is character-
ized by outwash (stratified drift). The iMars Hill Conglomerate
derived from Mars Hi 11 comprises many of the clasts in this pit.

Turn right out of the pit and farm. The hill to the left is in
New Brunswick.

61.2 Turn right at junction. Travel north along Mars Hill noting Green
Mountain Ridge to the left. As you approach Mars Hill bedrock
approaches the surface.

64.0 Turn left and follow dirt road.

64. 9 Continue on asphalt road.

65.6 Turn left on route 1A and head for the town of ' Mars Hill. Note
that the topography begins to level out toward the south.

66.6 1A joins route 1 south. Continue down route 1 to route 95. You are
now entering a pitted outwash plain. From Mars Hill to route 95 is
a dangerous road with frequent accidents involving trucks. Be aware

93.7 Take right onto 95 south at Houlton.
132.0 Take Sherman exit off 95 south.

132.3 Turn left on route 158 toward Sherman Mills. Striation locality.

133.8 Bear right beyond Gulf station. Stay on route 158 heading toward
Macwahoc.

136.6 Take right on route 2.

142.7 Take left by old shingle shack off route 2. This dirt logging road
is designated 06-20-10.

145.6 Take right onto dirt road 6.

148.4 Bridge over the Lower Macwahoc River. Park the cars along the
road. Be certain that your car is well off the road as this is
a main logging road with the trucks really rolling.

Stop 5. The Lower Macwahoc Till. This is a weathered, silty till
very similiar in texture to the St. Francis Till, which under-
lies Van Buren Till, at localities along the St. John River. This
stop is approximately 2.7 miles southeast of the Lower Macwahoc
Moraine, which will be our next locality.

Return to cars and continue on road 6.

148.9 Junction with road 6 and 7. Take left onto 7.

261

150.0 Junction with road 7 and 8. Take left onto 8.
150.3 Junction with road 8 and 8.1. Continue on road 8.

150.8 Junction with road 8 and 8.2. Take a left onto road 8.2.

151.2 St^E-^J. Lower Macwahoc Moraine. This moraine was exposed during
salvage operations of a spruce budworm infested forest. The
moraine has been traced for approximately two miles before
becoming obscured by dense vegetation. The fabric of the im-
brecated clasts indicated lodgement by ice flowing south.

Return to cars and return to route 2.

160.2 Take left onto to route 2 and head toward Macwahoc.

167.6 Molunkus Stream picnic area. The road now follows an esker dipping
on and off for a short distance.

169.9 At Molunkus take right on route 2.

178.7 Mattawamkeag. Take right on route 157 and head toward Millin-
ocket .

186.9 Crossing the Salmon River.

187.3 Take left into cut of the Penobscot Esker.

Stop 1 L Penobscot Esker gravel pit. This locality is at the approx-
imate latitude where the esker systems of northern Maine begin.
To the north, only small and insignificant esker forms are encoun-
tered.

Return to cars and head back to road.

188.7 Take left onto route 151 upon leaving the pit.

190.4 Junction with route 95 Interstate.

End of Trip

262

Trip C-7

STRUCTURE AND SEDIMENTOLOGY OF SILURO-DEVONIAN
BETWEEN EDMUNDSTON AND GRAND FALLS, NEW BRUNSWICK

P. Stringer and R.K. Pickerill
University of New Brunswick, Fredericton, N.B.

Introduction

The Ordovician, Silurian and Lower Devonian sediments of the
Edmundston - Grand Falls area in northwestern New Brunswick are exposed in
Acadian (Devonian) anticlinoria and synclinoria (Fig. 1) within which
subsidiary major and mesoscopic folds are present (Hamilton-Smith, 1970;
St. Peter, 1977). The anticlinorium near Grand Falls (Fig. 1) is part of
the Aroostook-Matapedia anticlinorium which extends from Maine across New
Brunswick to the eastern end of the Gaspe peninsula in Quebec (Rodgers,
1970; Williams, 1978). The lithologies of the folded formations and their
stratigraphic correlation within the area are indicated in Table 1. The
sediments are predominantly fine-grained greywackes and constitute the
Aroostook-Matapedia belt.

Tight upright folds and an associated steep to vertical cleavage are
characteristic structural features of the belt. As noted by Rodgers (1970,
p. 131), the folding must represent a very great shortening of the original
basin across the strike such that the present width of the belt is only a
relatively small fraction of its original width. No sign of the floor on
which the Ordovician to Lower Devonian sediments were laid down is visible
anywhere in the belt, and Rodgers (1970) suggested that if the floor was
shortened with the sedimentary cover it must have been non-sialic and
presumably was disposed of downward [by subduction] . Ordovician ocean
crust of probable Lower Ordovician age has been recognized, however, as
basement to the folded and cleaved Silurian-Lower Devonian strata of the
Chaleur Bay synclinorium southeast of the Aroostook-Matapedia anticlinorium
in northern New Brunswick (Stringer, 1975; Rast et alti 1976; Pajari
et at. j 1977; Rast and Stringer, 1980). Shortening of the cover rocks by
decollement on the underlying basement has been proposed (Rodgers, 1970,
p. 132; Stringer, 1975).

The folding and cleavage of Silurian-Lower Devonian rocks in the
Appalachian/Caledonian orogen have been attributed to late Silurian and
Devonian deformation resulting from continental collision during closure of
the Proto-Atlantic (Iapetus) Ocean (Dewey and Kidd, 1974), and in southern
Scotland the deformation, has been related directly to subduction during
the Silurian (Leggett et at., 1979; Stringer and Treagus, 1980) along the
northwest margin of the Iapetus Ocean (Cocks et at., 1980). In the
Canadian Appalachians, however, Williams (1979) has argued that the Iapetus
Ocean closed during the Taconian (Ordovician) orogeny and that the Acadian
orogeny was intraplate, subsequent to the closure of the Iapetus Ocean. In
New England, Robinson and Hall (1980) have suggested that the Acadian
orogeny was a result of Siluro-Devonian convergence between adjoining
continental plates, because closure of an ocean cannot be proven and it is
possible that the tectonic features formed through development of a west-

263

Figure 1.

Generalized geological map of the Edmundston - Grand falls area,
New Brunswick (based on St. Peter, 1977, fig. 2). Inset map
shows Acadian (Devonian) anticlinoria and synclinoria in New
Brunswick and adjacent Maine and Quebec (based on Rast and
Stringer, 1974), the Carboniferous cover (ornament), and the
location of the trip area.

dipping subduction zone within a former single continental plate. Whatever
the cause, the folding and cleavage attest to very significant shortening
during the Acadian (northern Appalachians) and Erian (Ireland to
Scandinavia) orogeny. In the Edmundston - Grand Falls area, Middle to
Upper Ordovician as well as Silurian-Lower Devonian sediments were
involved in the Acadian orogeny. Taconian deformation is lacking in the
Carys Mills Formation; some folding in the underlying unnamed unit may be
Taconian (Hamilton-Smith, 1969, 1970).

264

SERIES

Siegas

AROOST. - MAT.
St. Pe

ANTICLINORIUM
Fir . 1977

Hamilton - Smith, 1970

EAST

WEST

z
<

z
o

>

Ul
O

tc

Ul

o

-1

c

g

"5

E

Ul

Temiscouata
Formation

siltstone, slate,
arkose, greywacke,
conglomerate

•

•
in

•

•

z
<

S

D

-J

in

g

9
-1

Perham Formation

c
o

a

E
u

£

E
a

X
fa
•

a.

• •

Upper Member

slate, sandstone,
lithic wacke,
limestone

shale, si It stone, minor
sandstone

U

o

i

Lower Member
calcareous slate, minor
siltstone

Upsalquitch Fm.

siltstone , sandstone , 3
greywacke

u
•
>
o

C
O
-1

Sitgas Formation
sandstone, slatt, limestone

Matapedia Group
limestone, shale

Z
<

u

>
o

Q

a.
o

Carys Mills Fm.
limestone, slatt

Unnamed Unit

slatt, shale , sandstone

Grog Brook Group

argillite, sandstone, greywacke;
minor conglomerate , limestone

Table 1 Lithology and stratigraphic correlation of Ordovician to Lower
Devonain formations in the Edmunston-Grand Falls area, New
Brunswick.

265

Folding and cleavage

The tight upright Acadian folds in the Middle Ordovician-Lower
Devonian sediments of the Edmundston - Grand Falls area trend NE-SW and
are similar in style to Acadian folds throughout much of New Brunswick
and adjacent regions of the northern Appalachians. Sporadic to well
developed cleavage associated with the folds is persistently steep or
vertical. The folds and cleavage have been designated as F- and S , and
the cleavage has been described as axial planar to the folds (St. Peter,
1977). Stereographic plots of bedding and cleavage (Hamilton-Smith, 1970,
figs. 10-14) suggest, however, that the cleavage strikes a few degrees
clockwise of the strike of the fold axial surfaces. Cleavage noncoplanar
with fold axial surfaces is a widespread feature of Acadian deformation in
the northern Appalachians (Stringer, 1975).

The S- cleavage, recorded as a slaty or fracture cleavage by St. Peter
(1977), is formed by parallel or anastomosing partings, spaced at about
10 mm intervals in sandstone beds and much more closely in siltstone and
muds tone (shale) beds. In thin sections of siltstones and mudstones at the
hinges of F- folds, the mesoscopic S_ cleavage partings are seen as dark
films (e. 0.005 mm thick) made of fine-grained white mica, opaque minerals
and irresolvable material, in which the micas are aligned subparallel or
parallel to the cleavage direction. In the microlithons (a. 0.04 mm thick)
between the cleavage films, fine-grained white mica, quartz and calcite
grains commonly show no obvious preferred orientation, apart from
occasional parallelism with bedding. Diagenetic chlorite grains within the
microlithons often display mineral cleavage planes oriented perpendicular or
at a high angle to the S- cleavage films, and muscovite laths are
occasionally intergrown with the chlorite grains parallel to the mineral
cleavage of the chlorite host. Compression within the microlithons
perpendicular to the S. cleavage is indicated by crenulation of the chlorite
mineral cleavages and the intergrown muscovite laths. Where thin
calcareous or quartzose siltstone beds are present within the folded
mudstones, the S- cleavage films converge towards re-entrants in the silty
layers formed by buckles and by reverse microfaults (Figs. 2 and 4). The S
cleavage films correspond to pressure solution planes (Durney, 1972;
Williams, 1972; Gray, 1979; Stephens et at., 1979), and the cleavage
should therefore be described as a pressure solution cleavage. The
pressure solution cleavage films represent a significant component of the
shortening normal to the cleavage direction (Gray, 1979, pp. 114-7).

Evidence for the origin of the S cleavage films by pressure solution
is provided by a progressive development of microstructures (Gray, 1979,
fig. 3) which can be recognized in thin sections of the S. cleavage from
the hinges of F folds in laminated silty mudstones in the Edmundston -
Grand Falls area (Fig. 3) . Shortening in the mudstone is represented by
uniformly and closely spaced pressure solution S, cleavage films, and by
compression of the intervening microlithons indicated by the crenulated
chlorite grains. Shortening in the interbedded thin silty layers, in
which pressure solution films commonly are lacking, occurred by symmetric
to asymmetric buckling concomitant with development of the pressure
solution films in the mudstone. Adjacent asymmetric buckle folds may

266

o

in

•9
00

o
z

p

B

H

to

i

s

V)

o

to

s

a) co

> .e

c >-•

O (U

CJ P-!

co c

Tj cd

0) -H

tt) «H

C «H

O CO
4J
CO

T3 •

3 CO

6 T3

0)

a .o

IM

a)
C
o

■u
CO
4-1

CO «H

bo co

> C

CO tH

0) .C

rH 4-1
O

a

O *~s

•h co m
4-i »a

3H ft

.H O O

O «4-i 4-1

CO O CO

<u u

0 CJ

3
CO

CO o
CU «H

u u

P- 4-1
CU

5

CO

c

PQ

co e

T3 CO 0)

cu rt 2;

C

CU
CJ

O T3

u

CO CO

4.) 6 o

•h cu

T3 iH hJ

6

CO 4-» •

U 4-1

•a O CO

CU .C
CO

cO

u

4-i

c
cu
o
C
O

u

CM

CU
U

00
•H

CU

o

•H

4-t

CO

c
o

Pu O

3 fc

267

(a)

(c)

(d)

Figure 3.

Schematic diagram illustrating the progressive development
of microstructures in laminated silty mudstone during pressure
solution on S cleavage planes in the mudstone concomitant
with asymmetric buckling of a thin silty layer. (a) Evenly
spaced S pressure solution films in the mudstone show slight
concentration and thickening adjacent to short limb of gentle
asymmetric microfold in the silty layer. (b) and (c)
Progressive concentration, thickening and convergence of S-
pressure solution films against the rotating and partly
dissolved silty layer in the short limb, and development of
a second asymmetric microfold with opposite vergence. (d)
Complete solution of short limb produces apparent offset of
silty layer along thickly concentrated S pressure solution
films .

SCALE IN

Figure 4.

Divergent pressure solution films
associated with reverse micro-
faulting of a siltstone bed are
superimposed upon uniformly spaced
S pressure solution films in the
mudstone bed.

268

show opposing vergence. Pressure solution films in the muds tone are more
closely spaced and slightly thicker near the short limbs of the asymmetric
folds, indicating more intense solution of the microlithons (Fig. 3a).
With progressive rotation of the short limbs, the pressure solution films
in the mudstone become more concentrated, converging towards each side of
the short limb (Fig. 3b), and the silty layer is progressively dissolved
(Fig. 3c). The final stage is marked by loss of the short limb and the
development of a continuous zone of concentrated pressure solution films
along which the silty layer appears offset, resembling a microfault (Fig.
3d) . The different stages of microstructural development can often be
traced across several silty layers within a single thin section, from gentle
asymmetric microfold to apparent microfault.

Shortening of the thin silty layers by actual microfaulting is
indicated in places by low to high angle reverse microfaults, commonly
associated with the short limbs of asymmetric microfolds. Concentrations of
pressure solution films which diverge away from the reverse fault in the
silty layer into the adjacent mudstone may be superimposed upon uniformly
spaced pressure solution films formed in the mudstone prior to the micro-
faulting (Fig. 4) . Thus polyphase pressure solution microstructures may be
produced during a single phase of shortening.

Slump folding

Recumbent folds which pre-date the F folding and the S cleavage are
common in the Ordovician-Lower Devonian sediments of the Edmundston -
Grand Falls area (Stops 2 and 6). Acadian folding of the earlier folds is
prominent in the Lower Devonian Temiscouata Formation (Stop 3) , and has
been described in the Lower Silurian Siegas Formation (Hamilton-Smith, 1970,
pp. 29-30). The earlier folds exhibit no constancy in scale, style or
orientation, and lack axial plane cleavage (St. Peter, 1977), and have been
interpreted as slump folds in unconsolidated water-laden sediments. Large
scale recumbent pre-F folds indicated by inverted bedding in the Upper
Ordovician-Lower Silurian Carys Mills Formation in the Woodstock area (Rast
cum at, , 1980) 110 km south of Grand Falls have been interpreted by Rast
as soft-sediment deformation features due to submarine slip, although
Rast noted (ibid., p. 7) that it is very strange that no soft sediment
brecciation is associated with the large scale structures.

Sedimentology

Apart from the absence of late Silurian to early Devonian (Pridolian
to Gedinnian) strata, which St. Peter (1977) interprets as a Ludlovian to
Gedinnian disconformity, sedimentation within the area between Edmundston
and Grand Falls was essentially continuous from at least Middle Ordovician
(zone of Nemagraptus gracilis) to Lower Devonian (Emsian) time. This
sedimentation occurred within a narrow northeast/southwest elongate
trough extending from Gaspe through New Brunswick and into Maine and termed
the Central Clastic Belt and/or Aroostook-Matapedia Carbonate Belt by
Ayrton et at* (1969). Strata in this belt have been assigned various
group and /or format ional names at different places along the belt.

269

Within the Edmundston and Grand Falls section of the belt, relatively
few and detailed sedimentological analyses have thus far been undertaken,
notable exceptions being the studies of Hamilton-Smith (1970, 1971a, 1971b)
in the Siegas area and St. Peter (1977) in the Edmunds ton-Grand River area.
It is apparent from these studies that the Middle Ordovician-Lower
Devonian strata were deposited in a deep-water trough, within which a series
of isolated uplifted areas provided sedimentary detritus. Evidence of
deep-water deposition is not only provided by detailed analysis of
sedimentary lithofacies and depositional mechanisms but also by the
existence of a deep-water Nereites ichnofacies recently described by
Pickerill (1980) immediately to the northeast in the Matapedia district
in northern New Brunswick.

Sedimentological aspects emphasized on this excursion will be
directed essentially toward the Late Ordovician-early Silurian Carys Mills
(Stop 6) and the early Llandovery Siegas (Stop 4) Formations (Table 1).
The Siegas Formation is in part clearly erosive into the underlying Carys
Mills Formation, the estimated depth of erosion being in the order of a.
25-50 m (Hamilton-Smith, 1971b). Proximal turbidites with partial and/or
more complete Bouma sequences and we 11- developed resedimented limestone
conglomerates characterize these erosive fills. Associated massive and
structureless sandstone beds most probably represent deposition from grain
and/or fluidized flows. Proximal-distal relationships and palaeocurrent
analysis undertaken by Hamilton- Smith {ibid,) suggest derivation from the
north and north-west from an isolated topographic high. The Siegas
Formation was thus a localized channel sequence developed within the Carys
Mills Formation. More distal lithofacies of the former are undoubtedly
lateral facies equivalents of the latter.

The Carys Mills Formation and lateral equivalents throughout the
belt {e.g. Smyrna Mills Formation, Matapedia Group, eto.) consist
essentially of thinly interbedded calcareous argillites and argillaceous
calcitic and ankeritic limestones with thin and randomly developed graded
calcarenites with a substantial quartz component. Differential solution
has commonly etched the more limy beds to a greater extent so that the
rocks have a markedly ribbed appearance and are commonly referred to as
'ribbon limestones' (Ayrton et al., 1969). Previous authors {e*g. Ayrton
et at., 1969; St. Peter, 1977) have interpreted the Carys Mills Formation
as having resulted from deposition by turbidity currents in a distal
palaeoenvironment. We believe, however, that this is an oversimplification
and that whilst some units are suggestive of deposition from turbidity
currents, the majority of units are not. The thinly bedded calcarenites
exhibit partial Bouma sequences, particularly T and T , sequences, are
often erosive and contain broken and abraded shallow-water benthic marine
fossils. We believe these units to be the result of deposition by
turbidity currents. The thinly bedded calcareous argillites and
argillaceous calcitic and ankeritic limestones are, however, considered to
result from deposition by hemipelagic and normal bottom-following contour
currents viz:- contourites. Though the differentiation of contourites and
turbidites in the geologic record is extremely difficult and often
hazardous because of the lack of suitable distinguishing criteria (Bouma
and Hollister, 1973), the following observations are suggestive of

270

deposition by deep thermohaline boundary (contour) currents which
paralleled the ancient palaeoslope.

(i) The presence of abundant intraformational slump horizons
within the Carys Mills Formation together with minor resedimented debris
flow and olistostromal horizons. Such an association is more consistent
with slope (both modern and ancient) environments rather than distal or
overbank turbidite associations .

(ii) The widespread occurrence of similar facies which parallel the
presently exposed margins of the trough.

(iii) The presence of distinctive submarine channel lithofacies
which cut the regional strike and palaeocurrent (e.g. the Siegas
Formation) of the Carys Mills Formation.

(iv) The presence of a trace-fossil assemblage more consistent with
slope deposition rather than deep bathyal or abyssal palaeoenvironmental
conditions (e.g. amongst others, vertical burrows which cross-cut
successive units, Dipliehnites, Fuousopsis, Gyrochorte and spreiten-
bearing forms) .

(v) The presence of lenticular and irregular, commonly bioturbated,
horizons comparable to the 'muddy contourites' of Stow and Lovell (1979).

It is therefore suggested that the Carys Mills Formation represents
a succession of contourites and hemipelagic deposits which include
randomly and thinly developed turbidites. In modern regimes this
association is most consistently developed in continental rise environments
(Stow and Lovell, 1979). Nevertheless, as outlined above, evidence in the
Carys Mills Formation suggests that such an association may be equally as
common in an ancient slope environment.

References

Ayrton, W.G. , Berry, W.B.N. , Boucot, A.J., Lajoie, J., Lesperance, P.J.,
Pavlides, L. and Skidmore, W.B. 1969. Lower Llandovery of the
northern Appalachians and adjacent regions. Geol. Soc. Amer. Bull. ,
80, pp. 459-484.

Bouma, A.H. and Hollister, CD. 1973. Deep ocean basin sedimentation.
In: G.V. Middleton and A.H. Bouma (Editors), Turbidites and Deep
Water Sedimentation. Soc. Econ. Paleontol. Min. , Tulsa, Okla. , pp.
79-118.

Cocks, L.R.M., McKerrow, W.S. and Leggett, J.K. 1980. Silurian

palaeogeography on the margins of the Iapetus Ocean in the British
Isles. In: D.R. Wones (Editor), the Caledonides in the U.S.A.
Virginia Polytechnic Inst, and State Univ., Mem. 2, Blacksburg, Va. ,
pp. 49-55.

271

Dewey, J.F. and Kidd, W.S.F. 1974. Continental collisions in the

Appalachian-Caledonian orogenic belt: variations related to complete
and incomplete suturing. Geology, 2, pp. 543-546.

Durney, D.W. 1972. Solution-transfer, an important geological deformation
mechanism. Nature, Lond. , 235, pp. 315-317.

Gray, D.R. 1979. Micros tructure of crenulation cleavages: an indicator
of cleavage origin. Am. J. Sci. , 279, pp. 97-128.

Hamilton-Smith, T. 1969. Sedimentation during the Taconic orogeny: a

study of late Ordovician and early Silurian rocks of the Siegas area,
New Brunswick. Unpublished S.M. and S.B. thesis, Massachusetts
Institute of Technology.

Hamilton-Smith, T. 1970. Stratigraphy and structure of Ordovician and
Silurian rocks of the Siegas area, New Brunswick. Report of
Investigation No. 12, Mineral Res. Branch, Dept. of Nat. Resources,
New Brunswick, 55p.

Hamilton-Smith, T. 1971a. Paleogeography of northwestern New Brunswick
during the Llandovery: a study of the provenance of the Siegas
Formation. Can. J. Earth Sci., 8, pp. 196-203.

Hamilton-Smith, T. 1971b. A proximal-distal turbidite sequence and a
probable submarine canyon in the Siegas Formation (early
Llandovery) of northwestern New Brunswick. J. Sed. Petrol. , 41, pp.
752-762.

Legget, J.K., McKerrow, W.S. and Eales, M.H. 1979. The Southern Uplands
of Scotland: a Lower Palaeozoic accretionary prism. J. Geol. Soc.
Lond., 136, pp. 755-770.

Pajari, G.E., Rast, N. and Stringer, P. 1977. Paleozoic vulcanicity along
the Bathurst-Dalhousie geotraverse, New Brunswick, and its relations
to structure. In: W.R.A. Baragar, L.C. Coleman and J.M. Hall
(Editors), Volcanic Regimes in Canada. Geol. Assoc. Can., Spec. Paper
16, pp. 111-124.

Pavlides, L., Mencher, E. , Naylor, R.S. and Boucot, A.J. 1964. Outline of
the stratigraphic and tectonic features of northeastern Maine. U.S.
Geol. Surv. , Prof. Paper 501-C , pp. C28-C38.

Pickerill, R.K. 1980. Phanerozoic flysch trace fossil diversity -
observations based on an Ordovician flysch ichnofauna from the
Aroostook-Matapedia Carbonate Belt of northern New Brunswick. Can.
J. Earth Sci., 1_7 (in press).

Rast, N. and Stringer, P. 1974. Recent advances and the interpretation
of geological structure of New Brunswick. Geosci. Can., 1 (4), pp.
15-25.

272

Rast, N. , Stringer, P. and Burke, K.B.S. 1976. Profiles across the

northern Appalachians of Maritime Canada. In: C.L. Drake (Editor),

Geodynamics: Progress and Prospects. Am. Geophys. Union, Washington,
D.C. , pp. 193-202.

Rast, N. with St. Julien, P., Stringer, P., Pickerill, R.K. , Grant, R.H.

and Keppie, J.D. 1980. The northern Appalachian geotraverse: Quebec -
New Brunswick - Nova Scotia. Geol. Assoc. Can./Min. Assoc. Can.,
Field Trip Guidebook, Trip 3.

Rast, N. and Stringer, P. 1980. A geotraverse across a deformed Ordovician
ophiolite and its Silurian cover, northern New Brunswick. Tecton-
ophysics (in press) .

Robinson, P. and Hall, L.M. 1980. Tectonic synthesis of New England. In:
D.R. Wones (Editor), the Caledonides in the U.S.A. Virginia
Polytechnic Inst, and State Univ., Mem. 2, Blacksburg, Va. , pp. 73-82.

Rodgers, J. 1970. The Tectonics of the Appalachians. Wiley Intersci. ,
New York, 271p.

St. Peter, C. 1977. Geology of parts of Restigouche, Victoria and

Madawaska counties, northwestern New Brunswick. Report of

Investigation No. 17, Mineral Res. Branch, Dept. of Nat. Resources,
New Brunswick, 69p.

Stephens, M.B., Glasson, M.J. and Keays, R.R. 1979. Structural and
chemical aspects of metamorphic layering in metasediments from
Clunes, Australia. Am. J. Sci., 279, pp. 129-160.

Stow, D.A.V. and Lovell, J.P.B. 1979. Contourites: Their recognition
in modern and ancient sediments. Earth-Sci. Reviews, 14, pp. 251-
291.

Stringer, P. 1975. Acadian slaty cleavage noncoplanar with fold axial
surfaces in the northern Appalachians. Can. J. Earth Sci. , 16,
pp. 949-961.

Stringer, P. and Treagus, J.E. 1980. Asymmetrical folding in the Hawick
Rocks of the Galloway area, Southern Uplands, Scotland. Scott. J.
Geol. (paper submitted) .

Williams, H. (compiler). 1978. Tectonic-Lithofacies Map of the

Appalachian Orogen. Memorial University of Newfoundland, Map No. 1.

Williams, H. 1979. Appalachian Orogen in Canada. Can. J. Earth Sci.,
16, pp. 792-807.

Williams, P.F. 1972. Development of metamorphic layering and cleavage

in low grade metamorphic rocks at Bermagui, Australia. Am. J. Sci.,
272, pp. 1-47.

273

Itinerary

Assembly point is in the parking lot of Keddy's Motor
Inn, Presque Isle. Starting time 8:00 AM (Eastern
Standard Time). Upon leaving the parking lot, drive out
of Presque Isle and follow Route 1 for 60 miles (96 km)
north via Van Buren and Madawaska to Edmundston. Proceed
through Canada Customs and Immigration at the international
border between Madawaska, Maine and Edmundston, New Brunswick,
and drive 2 miles (3.2 km) north through Edmundston East to
the Edmundston Motel on Route 2 (Trans-Canada Highway).

Mileage Km

00.0 00.0 Stop 1. Park on southwest side of Trans-Canada Highway

(Route 2) at front of Edmundston Motel, Edmundston. Cross
highway to 100 m long road section.

Lower Devonian Temiscouata Formation. Dark grey laminated
shaly siltstones are deformed by symmetrical NE-SW trending
Acadian open folds (half-wavelength 50 m) and subveftical S,
pressure solution cleavage striking 052°. Cleavage /bedding
intersections plunge NE between 10° and 40°. At the
northwest end of the road section, small scale (0.5 cm) load
casts in thinly bedded fine grained sandstone interbeds in
the hinge of a syncline indicate that beds are the right
way up. About half way along the road section,
spaced composition banding parallel to vertical cleavage in
gently inclined beds indicates differential mineral
concentration along cleavage planes by pressure solution.

Return to vehicles and continue east on Route 2.

2.3 3.7 Stop 2. Park on south side of highway at roadside outcrop
30 m west of overhead power transmission lines.

Slump structures in Temiscouata Formation. At the west end
of the outcrop, tight minor recumbent slump folds are
exposed at the top of the rock faces within gently inclined
bedding. At the east end of the outcrop, aim thick
horizon of irregular slump structures dips SE within the
limb of an upright NE-SW trending Acadian fold with sub-
vertical S pressure solution cleavage which strikes 053°
and cuts through the slump structures. In places, a weak
second cleavage, S«, is present at a slight angle to the
S_ cleavage.

Return to vehicles and continue east on Route 2.

5.0 8.0 St. Basile

7.3 11.7 Stop 3. Park on south side of highway by the Weigh Station
sign. Cross highway to outcrop on north side of road.

274

Mileage Km

Folded larger scale recumbent slump fold, Temiscouata
Formation. A tight recumbent slump fold in dark grey
laminated shaly siltstone exposed for 20 m along its axial
surface is folded around the hinge of a NE-SW trending
Acadian anticline. S pressure solution cleavage inclined
steeply NW cuts through both limbs of the slump fold.
Sedimentary structures in the core of the anticline indicate
that bedding in the lower limb of the slump fold is
inverted. There is no fabric parallel to the axial surface
of the slump fold. Minor isoclinal slump folds verging
southeast are discernible on the upper limb of the larger
scale slump fold near the east end of the outcrop.

Return to vehicles and proceed east on Route 2.

11.3 18.2 Cross bridge over Green River (Riviere Verte) .

17.3 27.8 Ste. Anne de Madawaska

19.9 32.0 Turn left off Route 2, follow gravel road round back of
house and proceed east on old paved road.

20.2 32.5 Stop 4. Park on road near gates to Siegas Quarry. Walk
300 m north along road into quarry.

Lower Silurian Siegas Formation. The Siegas Quarry is
located on the northwest flank of the Aroostook-Matapedia
Carbonate Belt of Ayrton et al» (1969) , the most prominent
large structure of northeastern Maine and northwestern
New Brunswick during the Taconic orogeny (Hamilton-Smith,
1971a) . Although borderlands on both sides of this belt
were deformed, uplifted and eroded during the Taconic
orogeny, the Aroostook-Matapedia Carbonate Belt
continuously received sediment so that no unconformity of
any regional significance was developed within this belt
during Middle Ordovician-Late Silurian times. The charac-
teristic rocks of this part of the belt are known regionally
as the Carys Mills Formation, a succession of calcareous
thinly bedded flysch containing graptolites of Middle
Ordovician (Orthograptus truneatus var. intermedius) to
early middle Llandovery (Monograptus communis) age. The
Siegas Quarry, however, exposes rocks of a locally
developed clastic facies within this more regionally
developed calcareous flysch known as the Siegas Formation.
This formation ranges in thickness from c. 105-240 m in
the Siegas district and thins to zero within 10 km to the
south (Hamilton-Smith, 1971a).

The Siegas Quarry represents the principal reference section
of the Siegas Formation (Hamilton-Smith, 1970) and exhibits

275

Mileage Km

excellent exposures of limestone conglomerates, sandstones,
siltstones and shales. Fragmentary brachiopod genera include

Stricklandiaj Pleetothyrella> Leangella, Mendaeellcij
Eopleotodontdj Dalmanella and Protatrypa and attest to an
early Llandovery age. These and additional faunas are
listed in Ayrton et at, (1969).

The limestone-conglomerates and sandstones exhibit features
indicative of an origin by turbidity current grain support
mechanisms. These beds, which range in thickness from 4 cm -
8 m, show excellently developed clast imbrication, partial
and more complete Bouma sequences and a whole variety of sole
structures, such as gutter casts, prod, bounce and other tool
markings, flute casts, load casts, ripple marks and
dewatering features. Detailed palaeocurrent analysis by
Hamilton- Smith (1971a) indicated a provenance from the north
and northwest, which he suggested to be an isolated
relatively discrete uplift similar to others previously
described in northeastern Maine by Pavlides et at, (1964) .
The interbedded siltstones and shales are generally 1-100 cm
thick and exhibit parallel or more rarely cross-laminations.
They represent material reworked and/or deposited by normal
bottom marine currents. Organic activity is evidenced by the
obscure trails preserved on the upper surfaces of some
siltstones and shale beds as well as the more recognizable
ichnogenera Chondrites, Planolites3 Fuausopsis and
Dvpliohnites . Comparison of the Siegas Facies association
with similarly described ancient analogues suggests
deposition in a submarine channel or slope/base of slope
environment. The localized development of the formation
suggests that a channel is probably more realistic.

More detailed and complete descriptions of the Siegas
Formation lithofacies and regional relationships may be
obtained in Hamilton-Smith (1969, 1970, 1971a, 1971b).

Return to vehicles and continue east along old paved road.

20.5 33.0 Turn left onto Route 2 and continue eastward.

25.0 40.3 Stop 5. Park on south side of highway at roadside outcrop
600 m west of Route 2 /Route 17 interchange, St. Leonard.

Silurian Perham Formation. A NE-SW trending Acadian syncline
plunging gently SW at the east end of the outcrop deforms
slightly calcareous grey laminated silty mudstones. Intense
small-scale buckling of thin (1-5 mm) siltstone beds
indicates considerable shortening concomitant with pressure
solution on the S- cleavage planes in the mudstone beds (Fig.
2). The S cleavage dips steeply SE. Convergent cleavage
refraction is prominent in the thicker (3 cm) silty beds.

276

Mileage Km

Return to vehicles and continue east on Route 2.

32.6 52.5 Cross bridge over C.N.R. tracks.

36.1 58.1 Turn right off Route 2 at interchange exit for Grand Falls.

Cross bridge over Route 2 and follow road down to Grand Falls.

38.0 61.2 Stop 6. Turn right off road into parking area of Falls View
Motel, Grand Falls. Clamber down bank to extensive outcrop
at confluence of Saint John River and Little River.

Upper Ordovician-Lower Silurian Carys Mills Formation,
Aroostook-Matapedia anticlinorium. Beds striking NE-SW
and dipping steeply SE young towards the southeast and are
intersected by S pressure solution cleavage which strikes
ENE-WSW and dips 45-85° NNW. The e. 100 m thickness of
beds exposed includes numerous horizons of spectacular
intraformational slump folding. In places, the axes of
tight slump folds and convolute structures plunge steeply
SE at a maximum down the dip of the beds, but most of the
slump folds are variably oriented.

The majority of the beds exposed here are composed of the
more typical 'ribbon limestones' of the Carys Mills
Formation. The argillaceous calcitic and ankeritic
limestones are internally quite variable, some exhibiting
grading, some parallel lamination, some cross-lamination,
whilst others exhibit bioturbation or appear to be generally
structureless. The majority of beds are laterally
continuous whilst others are lenticular and irregular.
Associated thin and graded calcarenites may also be observed,
many of which have been completely dismembered due to in situ
soft-sediment deformation. Without doubt, however, the most
obvious soft-sediment deformation features are the spec-
tacularly developed intraformational slump fold horizons,
some of which are 2 m in thickness. At the southeastern
edge of the exposure occurs an enigmatic debris flow or
olistostromal horizon. The presence of this and the slump
folds attest to the presence of a slope during deposition
of the ribbon limestones.

Return to vehicles and drive out of motel car park,
turning right onto the road.

38.2 61.5 Turn right (south) following sign for Grand Falls Centre

and Fredericton, crossing bridge over the Saint John River.
Continue south along Broadway through Grand Falls centre.

38.8 62.4 At south end of Broadway turn left, following sign for

Route 2 East. Follow road up the hill out of Grand Falls.

277

Mileage Km

41.1 66.2 Turn left onto Route 2 at Grand Falls Portage, following

sign for Fredericton, Route 2 East. Proceed south on Route
2.

54.0 86.4 Stop 7 (optional). East side of highway (Route 2), north of
Four Falls.

Pre-Acadian (Silurian?) dykes in 'ribbon limestones' of the
Upper Ordovician-Lower Silurian Carys Mills Formation.

Return to vehicles and continue south on Route 2.

59.8 95.7 Turn right off Route 2 at Perth-Andover interchange onto
Route 19 (N.B.) /Route 167 (Maine) and drive west to Fort
Fairfield, passing through U.S. Immigration and Customs at
the international border. Follow Route 167 from Fort
Fairfield to Presque Isle. The distance back to Presque
Isle from Perth-Andover is 18 miles (29 km). End of field
trip.

278
TRIP D-l

STRATIGRAPHIC AND STRUCTURAL RELATIONS IN THE TURBIDITE
SEQUENCE OF SOUTH- CENTRAL MAINE

Philip H. Osberg
U.S. Geological Survey and University of Maine at Orono

Introduction

The turbidite section of south-central Maine occupies the southeast
part of the Merrimack synclinorium which extends from the vicinity of
Houlton, Maine, through southern Maine, central New Hampshire, central
Massachusetts, and into central Connecticut. These rocks belong to a thick
(>5000 m) section of Silurian age that is in contact to the east with
a terrane of plagioclase gneiss and amphibolite of Precambrian (?) or early
Paleozoic age. Early foliated plutons, presumably of pre-Silurian age,
intrude the gneiss-amphibolite terrane. Devonian plutons of granite, quartz
monzonite, and granodiorite cut both the turbidite and gneiss-amphibolite
terranes. The gneiss amphibolite terrane was metamorphosed to epidote-
amphibolite facies in Precambrian (?) or early Paleozoic time. Both the
turbidite and gneiss-amphibolite terranes were metamorphosed in a Buchan-
type facies series in Early Devonian time. The Buchan-type metamorphism
increases in grade to the south.

The distribution of rock units is shown in Figure 1. The region in-
cludes the Skowhegan, Pittsfield, Norridgewock, Waterville, Vassalboro,
Augusta, and Gardiner 15' quadrangles of the U.S. Geological Survey Atlas of
Topographic Maps. The geology is based on work of Barker (1961), Osberg
(1968), Heinonen (1971), Griffin (1973), Ludman (1976, 1977), Pankiwskyj ,
and others (1976), and on unpublished notes of Pankiwskyj, Newburg, and
Osberg.

Stratigraphy

The stratigraphic column is presented in the explanation of Figure 1.
The paleontological control comes from Osberg (1968) and from the data of
Pankiwskyj and others (1976).

The relative stratigraphic position of lithic units has been determined
by primary sedimentary features. But because of complex structural relations
and because both upward-facing and downward-facing sections have been recog-
nized, primary features are most relevant at or near formational contacts.
This circumstance limits the number of critical observations.

Cushing Formation. The Cushing Formation is exposed along the east boundary
of the area shown in Figure 1. These rocks are continuous southward to the
latitude of Portland, where the name Cushing was first used (Katz, 1917;
Hussey, 1971). The Cushing Formation in south-central Maine consists of
three divisions: (1) a lower quartz-K feldspar-plagioclase-biotite gneiss
consisting of quartzofeldspathic layers 1-3 cm thick alternating with
thinner biotite-rich layers, (2) a medial plagioclase-quartz-biotite gneiss

279

70°
45° "~~

EXPLANATION

69°05'

45°

Binary quartz monzonite

Biotite granodiorite

|\/\J Fall Brook Formation

Parkman Hill Formation

Unnamed phyllite/quartzite
unit

Sangerville Formation and
Limestone units

44°

Waterville Formation and
Limestone unit

Vassalboro Formation and
Limestone unit

Foliated granites

Wr*& Cushing Formation
KU Field Stop

44°

70'

A-Augusta, W-Waterville, S-Skowhegan

69°05'

Figure 1. Geologic map of south-central Maine,

280

interlayered with biotite-bearing amphibolites, and (3) an upper, somewhat
massive plagioclase-quartz-biotite granulite. A rusty zone, forming a
nearly continuous unit along the west boundary of the Cushing Formation,
may be a mineralized zone along a thrust contact or it may partly represent
a stratigraphic unit. All sequences contain abundant pegmatites. No thick-
ness can be given for the Cushing Formation.

Age relations for the Cushing Formation are not well established.
Mineralogy and texture from the north extremity of its outcrop indicate that
it has been metamorphosed to higher grade than the overlying Silurian turbi-
dite section. The Cushing Formation is probably early Paleozoic in age or
possibly Precambrian (?). The lithology and stratigraphic position of the
Cushing Formation are similar to those of parts of the "Massabessic" gneiss
in New Hampshire where Besancon and others (1977) have determined a 600 m.y.
age and to those of the Nashoba Formation in Massachusetts. The Cushing
Formation, however, cannot be traced continuously into these units.

Vassalboro Formation. The name Vassalboro sandstone was coined by Perkins
and Smith (1925) for massively bedded, bluish graywacke cropping out between
the Kennebec River and China Lake. The name has since been amended to
Vassalboro Formation (Fisher, 1941; Barker, 1961) because of the variability
of the unit. The name Kenduskeag Formation as used by Ludman and Griffin
(1974) has been applied to some rocks originally included within the
Vassalboro Formation and refers to a distinctive lithology, but it may not
represent a mappable unit.

The Vassalboro Formation Consists of bluish gray, slightly calcareous,
quartz wacke and quartz-mica phyllite/schist . The proportion of these
lithologies varies from place to place, but quartz wacke always makes up a
part of any outcrop. Thickness of bedding varies widely; the quartz wacke
is present in beds that range from 7 cm to several meters in thickness,
whereas the schist/phyllite is present in beds that range from a few milli-
meters to several centimeters in thickness. Beds of quartz wacke have
parallel lamination and, locally, cross lamination and convolute tops. At
higher metamorphic grades, the quartz wacke contains diopside, green amphi-
bole, clinozoisite, and abundant plagioclase. A prominent unit of quartz-
mica schist and gray marble observed at several localities near the east
boundary of the Vassalboro Formation may represent inliers of Waterville
Formation or a heretofore unrecognized stratigraphic unit within the Vassalboro,

The Vassalboro Formation has a large breadth of outcrop even though
attitudes of bedding are mostly vertical throughout the area. Consequently,
estimating the thickness of the Vassalboro is difficult. Its minimum
breadth of outcrop and its maximum thickness is approximately 3,200 m, but
presumably it is less than this figure due to folding within the unit.

A single locality containing identifiable fossils has been described
in the Vassalboro Formation (Pankiwskyj and others, 1976). Graptolites
from this locality suggest a Llandoverian to Ludlovian Age. Considering
the thickness of the Vassalboro Formation and the ages of the overlying
formations, I think that the Vassalboro is probably Llandoverian, although
the lower part of the formation could be as old as latest Ordovician.

The Vassalboro Formation is equated to the lower part of the Smyrna
Mills Formation of northeastern Maine, possibly part of the Berwick Forma-
tion in southern Maine, the Oakdale Quart zite and the Paxton Quartz Schist
in central Massachusetts, and the Hebron Formation in Connecticut. It may

281

correlate with the Quimby Formation in western Maine.

Waterville Formation. Perkins and Smith (1925) used the name Waterville
shale to designate rocks exposed west of the Kennebec River in the vicinity
of Waterville. Osberg (1968) changed the name to Waterville Formation,
which included an eastern and a western facies. Present usage restricts
the Waterville Formation to the eastern facies; the western facies is
included in the Sangerville Formation (Ludman, 1976).

The Waterville Formation is dominantly a thinly laminated phyllite/
schist. Typically, beds of quartzite or slightly calcareous quartzite
alternate with beds of quartz-mica phyllite/schist. Bedding thicknesses
are generally less than 3 cm. Convolute structures in the quartzitic layers
and grading between quartzitic and phyllitic layers can be observed locally.
The phyllite/schist layers contain a variety of mineral assemblages depend-
ing on metamorphic grade. Assemblages containing combinations of quartz,
plagioclase, muscovite, biotite, chlorite, garnet, staurolite, cordierite,
andalusite, and sillimanite are common.

A conspicuous unit composed of gray limestone and quartz-mica phyllite
forms an excellent mapping horizon within the Waterville Formation. This
unit is thinly bedded with limestone layers 1-6 cm thick alternating with
phyllite layers generally less than 4 cm thick. At high metamorphic grades,
this unit contains diopside, grossularite, green amphibole, phlogopite, and
plagioclase.

The thickness of the Waterville Formation cannot be determined from a
single section because of the complexity of folding. However, an estimate
of maximum thickness can be determined by measuring the narrowest breadth
of outcrop from its base to the limestone member and adding that value to
the minimum breadth of outcrop between the limestone member and the upper
boundary of the formation. The resulting figure is 1,100 m. Of course,
the thickness of the formation may vary geographically along and across
strike.

Three fossil localities in the Waterville Formation have been described
(Osberg, 1968; Pankiwskyj and others, 1976). Two of these localities con-
tain fossils that range in age between Ordovician and Early Devonian, but
the third contains fossils that limit the age to a range of late Llandover-
ian to Ludlovian.

The Waterville Formation may be correlative to the west with the
sequence Turner through Anasagunticook Formations as used by Pankiwskyj and
others (1976) and the Greenvale Cove Formation. It may correlate to the
south with at least part of the Eliot Formation.

Sangerville Formation. The Sangerville Formation includes exposures of
graywacke and phyllite in the Skowhegan quadrangle (Ludman and Griffin,
1974; Pankiwskyj and others, 1976). These rocks are in part continuous with
rocks that Osberg (1968) interpreted as the western facies of the Waterville
Formation, and the name, Sangerville Formation, is adopted for these rocks.
This unit also contains rocks that Osberg (1968) called the Mayflower Hill
Formation, a name which should now be abandoned.

282

The Sangerville Formation lies stratigraphically above the Waterville
Formation, and as used here consists of two principal lithologies: a lower
graywacke-phyllite unit and an upper feldspathic wacke and interbedded
limestone unit. The upper feldspathic wacke and interbedded limestone unit
makes up the bulk of the formation.

Slightly calcareous graywacke and quartz-mica phyllite are the common
rocks of the lower unit. Beds range from 7 cm to a meter in thickness, and
the thickness of the graywacke is eight or nine times that of the phyllite.
Beds are commonly graded. Parallel lamination is common and delicate cross-
beds can be seen locally. A black sulfidic quartz-mica phyllite is found
discontinuously at the contact with the Waterville Formation.

The upper unit consists principally of light-gray feldspathic graywacke
and medium-gray quartz-mica phyllite/schist . Some of the graywacke is
slightly calcareous. Beds are 8 cm to 1.5 m thick, and, although most beds
are ungraded, they locally show grading, cross lamination, and scour-and-
fill features. The graywacke is locally conglomeratic and contains clasts
of slate, feldspar, plutonic rocks and volcanic rocks. The phyllite/schist
is laminated to massive. At appropriate metamorphic grades, it contains
staurolite, garnet, and sillimanite.

Two conspicuous units of gray limestone are present in the upper unit.
Lithologically, the limestone is similar to the limestone member of the
Waterville Formation, a circumstance that has caused much confusion in
earlier studies (Osberg, 1968; Ludman and Griffin, 19 74, and Pankiwskyj
and others, 1976).

The thickness of the Sangerville Formation is difficult to estimate,
and the writer has found no good place to reconstruct its thickness.
Ludman (1977) gave a thickness in excess of 2,000 m.

Fossils are found at three localities in the area shown in Figure 1.

Graptolites from these localities indicate an age from late Llandoverian

to Wenlockian (Osberg, 1968; Pankiwskyj and others, 1976). This age range

is confirmed by fossil information from adjacent areas.

The Sangerville Formation is equivalent to the Rangeley Formation in
western Maine. No correlatives are known in southern Maine.

Unnamed phyllite/quartzite unit. Thin bedded greenish gray quartz-mica
phyllite and quartzite occupy the stratigraphic interval between the
Sangerville and Parkman Hill Formations. This unit was mapped as Waterville
Formation by Griffin (1973) and Pankiwskyj and others (1977) and was in-
cluded in the Sangerville by Ludman (1976) .

This unit consists of beds of quartzite from 1 to 8 cm thick alter-
nating with beds of quartz-mica phyllite/schist from 1 to 3 cm thick. The
quartzite commonly shows conspicuous crossbeds and small slump structures
near the tops of beds. The quartz-mica phyllite/schist contains garnet
and staurolite at appropriate metamorphic grades. This unit somewhat
resembles the Waterville Formation, but can be distinguished from it by the
thicker beds of quartzite and the common internal bedding features.

283

The unnamed phyllite/quartzite unit has a breadth of outcrop in the
vicinity of Skowhegan of approximately 150 m. Because of extensive folding,
a thickness of 50 to 70 m may be more realistic.

No fossils have been found in this unit. However, it lies between the
Sangerville Formation of Llandoverian to Wenlockian age and the Parkman Hill
Formation of probable Ludlovian age; therefore, its age must be Wenlockian
or Ludlovian.

This unit is lithologically and stratigraphically equivalent to the
Perry Mountain Formation in western Maine. It has no known correlatives in
southern Maine, New Hampshire, Massachusetts or Connecticut.

Parkman Hill Formation. Pankiwskyj and others (1976) described the Parkman
Hill Formation as consisting of sulfidic, rusty-weathering metasandstone and
metapelite that lies between the Sangerville and Fall Brook Formations. As
used in this report it refers to the rusty-weathered metasandstone and
metapelite that lies between the unnamed phyllite/quartzite unit described
above and the Fall Brook Formation.

As used here the Parkman Hill consists of rusty-weathering, pyritifer-
ous quartz-mica phyllite in beds 0.5 to 10 cm thick. Bedding is difficult
to discern on weathered surfaces. Beds of variously rusty, slightly
calcareous quartz wacke and graywacke are interbedded with the phyllite,
particularly in the western part of the area. The wacke beds have thick-
nesses up to 15 cm.

Pankiwskyj (1979) estimates a thickness of 200 to 300 m for the Parkman
Hill Formation in the Anson quadrangle, but in the vicinity of Canaan
(Skowhegan quadrangle) its thickness is only several tens of meters.

Numerous fossil localities have been found in the Parkman Hill Forma-
tion (Pankiwskyj and others, 1976). These fossils have been assigned to
the range late Llandoverian to Ludlovian, but the most definitive collec-
tions indicate a Ludlovian age.

The Parkman Hill Formation is correlative with the Smalls Falls
Formation of western Maine.

Fall Brook Formation. Pankiwskyj and others (1976) used the name Fall
Brook Formation for a quart zite sequence in central Maine. Ludman (1977)
extended the name to include rocks exposed east of Skowhegan (Figure 1).

The Fall Brook Formation consists dominantly of massively bedded (15 cm
to 4 m thick), slightly calcareous quartz wacke. Cross lamination is con-
spicuous in some beds, and calcareous pods are locally common. Amphibole,
diopside, and plagioclase are present at appropriate metamorphic grade.
Groups of beds, 1 to several meters thick, of interbedded quartz-mica
phyllite/ schist and quartzite in beds 1 to 3 cm thick form a subordinate
lithology .

No estimates of the thickness of the Fall Brook Formation can be made
within the area shown in Figure 1. Pankiwskyj and others (1976) state its
thickness to be 1,000 m.

284

The Fall Brook is unfossiliferous, but because it lies between the
rusty unit of Ludlovian age and Early Devonian phyllites, its age is Upper
Silurian or Lower Devonian.

The Fall Brook Formation correlates with the Madrid Formation in
western Maine. Its correlatives to the south have not been recognized.

Intrusive Rocks. The intrusive rocks are represented by conspicuously
foliated "older" granites, dikes of plagioclase granulite and stocks of
binary quartz monzonite and biotite granodiorite. The "older" granites are
represented by equigranular garnet and biotite bearing granites and pegma-
tites and by biotite granite containing insets of K feldspar. All varie-
ties are foliated. These "older" granites intrude the Cushing Formation,
but do not cut the Silurian turbidite section. These granites have not been
dated, but somewhat similar rocks to the northeast have a minimum zircon age
of 450 m.y.

Dikes of plagioclase granulite consist of plagioclase, calcite, musco-
vite, chlorite, and magnetite. These dikes are 60 cm to 3 m thick. They
crosscut upright isoclinal folds but are in turn deformed by younger folds.
They obviously predate the metamorphism.

Binary quartz monzonite is exposed in four small stocks and lesser
bodies (Figure 1). Plagioclase, microcline, quartz, muscovite, biotite,
and garnet form an interlocking, inequigranular texture.

Biotite granodiorite is present in large stocks (Figure 1). The
biotite granodiorite contains plagioclase, quartz, microcline, biotite,
and, locally, garnet or hornblende. It has a perceptible foliation, particu-
larly near the contacts.

The plagioclase dikes, binary quartz monzonite and biotite granodi-
orite cut Silurian rocks. The plagioclase dikes preceded metamorphism and
Ferry (1978) has presented arguments that the quartz monzonite and granodi-
orite were coeval with regional metamorphism. Several of the stocks have
been dated by the Rb/Sr method which indicates ages between 394 and 360 m.y.
(Dallmeyer, 1978).

Structural Geology

Structural features include schistosity, cleavage, a variety of
lineations, folds, cleavage bands, boudinage, shears, and faults. Fold
elements, assigned to three distinct episodes of deformation, consist of
late asymmetrical folds, earlier upright folds, and preexisting recumbent
folds.

The late asymmetrical folds have axial surfaces that are more or less
constant throughout the region (strikes within 10° of north and steep dips)
A well-developed cleavage parallels their axial surfaces. These folds
deform bedding and schistosity/cleavage of earlier folds; therefore, the
plunges and asymmetry of these late folds are variable and depend on the
orientation of the surface being folded. No major folds belonging to this

285

episode have been identified.

These late folds were approximately synchronous with the metamorphism
but the deformation continued after the peak of metamorphism. Muscovite
and sillimanite have nucleated in the surfaces of cleavage associated with
these folds, and porphyroblasts, although they exhibit static growth rela-
tive to earlier schistosity, have rotated slightly during the late deforma-
tion. Ferry (1978) suggested that the regional metamorphism and emplacement
of plutons of granodiorite and quartz monzonite were coeval; this relation-
tion would indicate that these folds are Early Devonian in age.

Upright folds are both mesocopic and major, and these folds control
the regional map patterns (Figure 1). In outcrop, these folds have forms
that range from relatively open folds to highly flattened isoclinal folds.
The openness depends on proximity to competent structural units. Axial
surfaces strike northeast and dip steeply; plunges are generally less than
25°, but locally plunges of 45° have been recorded. Regional schistosity
and pressure-solution cleavage parallel the axial surfaces of these folds.
These upright folds face downward over certain regions and upward over
others.

Major upright folds are the obvious structural features as indicated
from the distribution of formations (Figure 1). Major folds delineated in
Figure 1 are from northwest to southeast: the Athens antiform, the Corn-
ville antiform, the Skowhegan synform, the Hinckley antiform, and the
Shawmut antiform. Lesser antiforms and synforms are defined on the basis
of the limestone/marble members of the Waterville and Sangerville Forma-
tions.

The upright folds predate the metamorphism. The schistosity that is
parallel to their axial surfaces is statically enclosed in porphyroblasts,
and these folds are cut by dikes (Osberg, 1968) that have been metamor-
phosed to plagioclase granulites. On the other hand, these folds regionally
deform rocks as young as Oriskany and possibly as young as Schoharie.
Because the metamorphism has been dated at 400 to 365 million years (Faul
and others, 1963; Dallmeyer, 1978), these folds have an Early Devonian age
as well.

Evidence of the recumbent folds is mostly indirect: map-pattern,
downward-facing upright folds, and stratigraphic relationships. Only at a
single outcrop do minor folds display evidence of the earlier recumbent
folds. The reconstruction of these folds indicates that they are of Alpine
proportions and in south-central Maine face northwest.

The upright folds are interpreted to deform earlier recumbent struc-
tures. The Athens, Cornville, Skowhegan, and Hinckley upright folds face
upward and are thought to deform the normal limb of a large west-facing
recumbent anticline. Near Waterville the Shawmut antiform faces upward and
is interpreted to be in the normal limb of a west-facing recumbent sylcine.
The synform east of the Shawmut anticline and delineated by the limestone
member of the Waterville Formation faces downward.

286

• • .

# *

6

• •

CD

u

3
00
•H

O

CO

01
6
nj
co

CO

C
cu

I

oo

c
o
i— i

c
o

u
cu

CO
0)

4-1
O

3
U

4-1
CO

0)

u

3
00

tn

287

A high-angle fault separates the Waterville Formation from units to
the west except west of Waterville where different parts of the Sangerville
are juxtaposed (Figure 1). This fault postdates the upright folds and
places the Shawmut antiform on the east flank of the Hinckley antiform with-
out an intervening synform. This fault predates the metamorphism.

The contact between the Cushing and Vassalboro Formations is incom-
pletely known (Figure 1). It may be an unconformity or it may be a thrust
fault. In either case the contact is folded by the upright folds.

A structure section showing the relationships between the early
recumbent folds and the late upright folds is detailed in Figure 2. The
line of the structure section is shown in Figure 1.

Metamorphism

Metamorphic rocks of the Silurian turbidite section have been studied
by Osberg (1968, 1971, 1974, and unpub. data) and by Ferry (1976a, b). The
metamorphism belongs to a Buchan-type facies series. Osberg (1974) mapped
isograds in the pelitic rocks on the basis of biotite, garnet, andalusite-
cordierite-staurolite, sillimanite, and sillimanite-K feldspar. Ferry
(1976a) mapped isograds in the Vassalboro and Waterville Formations on the
basis of biotite-chlorite, amphibole-anorthite, zoisite, microcline-
amphibole, diopside, and scapolite.

Examination of the chemistry of the associated minerals indicates that
the mineral assemblages approached equilibrium near the peak of metamor-
phism (Osberg, 1971; Ferry, 1978). Estimates of the conditions of meta-
morphism (Osberg, 1971; Ferry, 1976b) from coexisting minerals indicate
that the temperature ranged from ~378°C at the biotite isograd to about 550°
at the sillimanite isograd. Pressures have been estimated to have been in
the range 3,000-3,500 bars. Water pressure was generally less than total
pressure.

Dallmeyer (1978) has dated biotite and hornblende from various meta-
morphic assemblages. Undisturbed ^Ar/^Ar reiease spectra and total-gas
determinations for biotite produce ages ranging from 335 to 227 m.y.; rocks
are younger in the direction of increasing metamoprhic grade. Hornblende
records older ages than does biotite, and together these ages must be
interpreted to reflect diachronous post-metamorphic cooling.

References

Barker, D.S., 1961, Geology of the Hallowell granite, Augusta, Maine:
Princeton, N.J., Princeton University, Doctoral thesis, 240 p.

Besancon, J.R., Gaudette, H.E., and Naylor, R.S., 1977, Age of the

Massabesic gneiss, southeastern New Hampshire [abs.]: Geol. Soc
America Abs. with Programs, 9(3), 242.

288

40 39
Dallmeyer, R.D., 1978, Ar/ Ar and Rb/Sr ages in west-central Maine:

Their bearing on the chronology of tectono-thermal events: Geol. Soc.

America Abs. with Programs, 10(2), 38.

Faul, Henry, Stern, T.W. , Thomas, H.H., and Elmore, P.L.D., 1963, Ages of
intrusion and metamorphism in the northern Appalachians: Amer. Jour.
Sci., 261(1), 1-19.

Ferry, J.M. , 1976a, Metamorphism of calcareous sediments in the Waterville-
Vassalboro area, south-central Maine: Mineral reactions and graphical
analysis: Amer. Jour. Sci., 276(7), 841-882.

, 1976b, P, T, f(X)2» anc* -^HoO during metamorphism of calcareous sedi-
ments in the Waterville-Vassalboro area, south-central Maine: Contrib.
Mineralogy and Petrology, 57(2), 119-142.

, 1978, Fluid interaction between granite and sediment metamorphism,

south-central Maine: Amer. Jour. Sci., 278, 1025-1056.

Fisher, L.W. , 1941, Structure and metamorphism of Lewiston, Maine, region:
Geol. Soc. America Bull., 52(1), 107-160.

Griffin, J.R. , 1973, A structural study of the Silurian metasediments of

central Maine: Riverside, Calif., University of California at River-
side, Doctoral Thesis, 157 p.

Heinonen, C.E., 1971, Stratigraphy, structural geology and metamorphism of
the Tacoma Lakes area, Maine: Orono, Maine, University of Maine at
Orono, Masters Thesis, 79 p.

Hess, P.C., 1969, The metamorphic paragenesis of cordierite in pelitic
rocks: Contrib. Mineralogy and Petrology, 24(3), 191-207.

Holdaway, M.J., and Lee, S.M. , 1977, Fe-Mg cordierite stability in high-
grade pelitic rocks based on experimental, theoretical, and natural
observations: Contr. Mineralogy and Petrology, 63, 175-198.

Hussey, A.M., II, 1971, Geologic map of the Portland quadrangle, Maine:
Maine Geol. Survey, Geol. Map GM-1, scale 1:62,500.

Katz, F.J., 19717, Stratigraphy in southwestern Maine and southeastern New
Hampshire: U.S. Geol. Survey Prof. Paper 108, p. 165-177.

Ludman, Allan, 1976, Fossil-based stratigraphy in the Merrimack synclinorium,
central Maine: in Page, L.R. (ed.) Contributions to the stratigraphy
of New England: Geol. Soc. America Memoir 148, 65-78.

, 1977, Geologic map of the Skowhegan quadrangle: Maine Geol. Survey

Geol. Map GM-5, scale 1:62,500.

289

Ludman, Allan, and Griffin, J.R. , 1974, Stratigraphy and structure of central
Maine: Jji Guidebook for field trips in east-central and north-central
Maine: New England Intercollegiate Geological Conference, 66th Annual
Meeting, Orono, Me., Oct. 12-13, 1974, Univ. of Maine at Orono, 154-
179.

Osberg, P.H. , 1968, Stratigraphy, structural geology, and metamorphism of
the Waterville-Vassalboro area, Maine: Maine Geol. Survey Bull. 20,
64 p.

, 1971, An equilibrium model for Buchan-type metamorphic rocks,

south-central Maine: Amer. Min. , 56, 569-576.
, 1974, Buchan-type metamorphism of the Waterville pelite, south-

central Maine: in Guidebook for field trips in east-central and north-
central Maine: New England Intercollegiate Geological Conference,
66th Annual Meeting, Orono, Me., Oct. 12-13, 1974, Univ. of Maine at
Orono, 210-222.

Osberg, P.H. , Moench, R.H. , and Warner, J., 1968, Stratigraphy of the

Merrimack synclinorium in west-central Maine: in Zen, E-an, and others
(eds.) Studies of Appalachian Geology; Northern and Maritime: New York,
Wiley- Interscience, 241-254.

Pankiwskyj , K. , 1979, Geologic maps of the Kingfield and Anson 15' quad-
rangles, Franklin and Somerset Counties, Maine: Maine Geol. Survey,
Geologic Map Series GM-7, 51 p.

Pankiwskyj, K. , Ludman, Allan, Griffin, J.R., and Berry, W.B.N. , 1976,
Stratigraphic relationships on the southeast limb of the Merrimack
synclinorium in central and west-central Maine: in Lyons, P.C., and
Brownlow, A.H. (eds.) Studies in New England geology: Geol. Soc.
America Memoir 146, 263-380.

Perkins, E.H., 1924, A new graptolite locality in centram Maine; with notes
on the graptolites by Rudolf Ruedemann: Amer. Jour. Sci. , 208, 223-227.

Perkins, E.H., and Smith, E.S.C., 1925, Contributions to the geology of

Maine, No. 1, A geologic section from the Kennebec River to Penobscot
Bay: Amer. Jour. Sci., 209, 204-229.

Thompson, A.B., 1976, Mineral reactions in pelitic rocks: II. Calcula-
tions of some P-T-X (Fe-Mg) phase relations: Amer. Jour. Sci., 276,
425-454.

Itinerary

Mileage

0 Assembly point for trip is Stafford's Variety in Athens, Maine
(11 miles N of Skowhegan) . Starting time is 9:00 A.M.

290

Stop 1. Outcrop is located on stream 300 m up stream from
bridge. Exposures of limestone member of Sangerville Formation.

Gray limestone and sandy limestone in beds 1-13 cm thick. Sandy
limestone beds exhibit well preserved cross bedding. Interbeds of
siliceous phyllite and biotite quartzite are 1-8 cm thick. Beds
strike northeasterly and dip steeply. Well preserved upright,
nearly isoclinal folds are preserved in bedding and the cross-
bedding indicates that these folds face upward. Isoclinal folds
plunge gently south.

Ludman (1977) mapped a fault that was interpreted to isolate this
limestone locality. Alternately, the map-pattern of this lime-
stone occurrence may be due to the intersection of the erosion
surface with a culmination in a major upright antiform.

Return to cars.

Proceed south on Route 150.

1.4 Townline between Athens and Cornville.

1.8 Cass Corner. Turn left on West Ridge Road.

6.0 Cornville. Continue straight. Hills to east are underlain by
biotite granodiorite of Hartland pluton.

8.5 Small outcrop of limestone member of Sangerville Formation to right.

8.7 Malbons Mills. Turn left.

8.8 Stop 2. Park along road. Outcrop is at old mill south of bridge.
Exposures of Sangerville Formation.

Under bridge is exposure of quartz-mica phyllite and quartzite in
beds 1-3 cm thick. Two thin beds of limestone are intercalated.

At old mill is exposure of massive feldspathic graywacke, poly-
mictic conglomerate, and quartz-mica phyllite. Bedding is 2 cm -
2 m thick. Clasts consist of vein quartz, wacke, feldspar, slate,
and volcanic fragments and tend to be flattened in plane of
bedding. Maximum size of clasts is approximately 1 cm. Bedding
and schistosity are essentially parallel; they strike northeast
and dip steeply west.

On east bank of stream exposures are feldspathic graywacke in beds
30-90 cm thick interbedded with 0.5-2.5 cm beds of quartz-mica
phyllite. Flame structures at bottom of sandy beds indicate tops
east. Late cleavage strikes more northerly than bedding and
schistosity and dips steeply west.

Return to cars and reverse direction.

8.9 Malbons Mills. Turn left toward Skohegan.

291

10.7 Stop sign. Junction with Route 2. Turn right onto Route 2.

10.8 Stop 3. Park along Route 2. Exposures of unnamed phyllite/
quartzite unit are on promontory west of Great Eddy of Kennebec
River. Route 2 is busy highway. Be cautious of traffic!

Quartzite beds 2-7 cm thick show well developed crossbedding.
Light gray quartz-mica phyllite, in beds 0.1-3 cm, alternate with
beds of quartzite. A few lenses of gray limestone, approximately
30 cm long and 5 cm thick are locally intercalated. Bedding
strikes northeast and dips steeply west; crossbeds indicate tops
are east. Nearly isoclinal upright folds deform bedding. Cross-
beds in the limbs of folds indicate that the folds face upwards.
A late cleavage strikes more northerly than bedding and dips
steeply northwest.

Return to cars. Reverse direction and proceed east on Route 2.

11.1 Stop 4. Turn right into "picnic area". Exposures of Parkman Hill
Formation are located on the bank of the Kennebec River. These
outcrops are in west limb of Skowhegan synform.

Rusty-weathering, dark gray pyritiferous quartz-mica phyllite in
beds 2-30 cm thick interstratif ied with beds of rusty-weathered,
light purplish gray pyritiferous quartz-feldspar-mica granulite,
8-90 cm thick. Bedding and schistosity strike northeasterly and
dip steeply west.

Return to cars and proceed east on Route 2.

12.6 Turn right onto East River Road.

14.5 Stop 5. Park along road. Small exposure of Fall Brook Formation
is on north side of road. Good exposures of Fall Brook Formation
are not accessible along our route.

Somewhat flinty, purplish gray quartz-plagioclase-biotite-calcite
granulite in beds 8-30 cm thick. Medium gray quartz-mica phyllite
in layers 1-3 cm thick is interbedded. Quartz veins are numerous.
These rocks lie well within the contact aureole of the Hartland
pluton. Bedding and schistosity strike northeast and dip steeply
west. A small right-hand fold deforms both bedding and a quartz
vein; it plunges moderately south. A late cleavage strikes more
northerly than bedding and dips steeply west.

Return to cars and proceed east on East River Road.

15.4 Turn left onto Eaton Mountain Road. Outcrops of rusty-weathering,
pyritiferous quartz-mica phyllite on left in brook belong to Park-
man Hill Formation and are in east limb of Skowhegan synform.

16.7 Eaton Mountain to right is underlain by Parkman Hill Formation and
biotite granodiorite of Hartland pluton.

292

18.0 Stop sign. Intersection with Route 2. Turn right and proceed
east on Route 2.

18.9 Small outcrop of Hart land granodiorite on left.

21.5 Canaan. Outcrop of limestone member of Sangerville Formation
under bridge.

21.8 Route 23 diverges. Continue east on Route 2.

22.9 Small exposures of Sangerville Formation on left.

23.2 Small outcrops of Sangerville Formation on right.

29.4 Junction of Route 152 with Route 2. Turn right on Route 152.

29.5 Stop 6. Park on side of road. Exposures of Waterville Formation.

Light purplish and grayish green quartz-mica-chlorite phyllite
containing 0.2-8 cm thick beds of quartzite and ankeritic quart-
zite. Bedding strikes northeasterly and dips steeply east.
Upright, isoclinal folds are present at north end of outcrop.
These folds plunge gently northeasterly. Foliation in the phyllite
wraps around the hinges of these folds.

Return to cars and proceed south on Route 152.

32.6 Entering Pittsfield. Stop sign. Turn right to Route 95.

33.4 Turn left onto southbound lane of Route 95.

35.7 Turn right into rest area. Lunch stop. Continue south on
Route 95.

44.1 Exposures of Waterville Formation on right.

47.5 Bridge over Kennebec River.

49.8 Exposures of rusty-weathering, sulfidic quartz-mica phyllite at
base of Sangerville Formation.

51.0 Bridge over Messalonski Stream.

53.1 Turn right off Route 95 using Exit 33.
53.4 Stop sign. Turn left onto Oakland Road.

54.3 Stop 7. Park at side of highway. Oakland Road is a busy street.
Be careful of traffic. Exposure of basal Sangerville Formation
is located on north side of highway.

Rusty weathered, dark-gray quartz-mica phyllite interbedded with
dark-gray quartzose layers at the east end of the outcrop is the
basal unit of the Sangerville Formation. Pyrite is abundant

293

and crystals have a more-or-less common orientation and do not
have associated, quartz-filled pressure shadows. Beds range from
5 mm to 1.5 cm in thickness.

Light-gray, slightly calcareous graywacke and quartz-mica phyllite
exposed in the west part of the outcrop are the dominant rocks of
the lower unit of the Sangerville Formation. Beds are 15 cm to
50 cm thick and commonly grade from sandy bottoms to phyllitic
tops. Sandstone: shale ratios in beds range from 1:1 to 8:1.
Small grains of feldspar and shale clasts can be seen at the base
of some beds. Good grading can be seen on top of the outcrop and
indicates that the cyclically graded unit is younger than the
sulfidic phyllite.

Bedding strikes northeasterly and dips steeply. Schistosity cuts
bedding at a low angle and the trace of graded beds on schistosity
indicate that the section is right-side up. Small shears that
offset bedding can be seen on the vertical face, but because the
topping direction of bedding does not change across them, the
displacement on them is thought to be small. Quartz pods form
boudin fillings; other veins are late.

The observed stratigraphic relations indicate that the Sangerville
Formation is younger than the Waterville Formation. The outcrop
pattern of the Sangerville Formation (Figure 1) in the Shawmut
antiform, coupled with the consistent northward plunge of the
Shawmut antiform, demands that the Shawmut antiform in part faces
downward, i.e., folds the axial surface of an earlier recumbent
fold. Because the local section is right-side up, this outcrop
must be in the normal limit of the early recumbent syncline
(Figure 2) .

Return to cars. Continue east on Oakland Road.

54.9 Stop light. Continue straight. After crossing bridge, bear
right on Grove Street.

55.4 Stop sign. Turn left on Water Street.

56.2 Stop light. Turn right onto Route 201 and cross bridge over
Kennebec River.

56.3 Stop light. Turn right remaining on Route 201.

56.5 Stop 8. Park in vacant lot to right. Exposure is located over
the bank on the Kennebec River. Outcrop is Waterville Formation
and its limestone member.

Alternations of light-gray quartz-muscovite-chlorite-phyllite
and white to buff quartzite in beds 6 mm to 8 cm thick is a common
variant to the Waterville Formation adjacent to the limestone
member. Some beds show gradation from quartzone bottoms to pelitic
tops. The limestone member consists of gray, slightly micaceous
limestone interbedded with rusty, buf f-collored quartz-mica.

294

phyllite. Bedding is 6 mm to 12 cm thick. The contact between
the two units is abrupt.

Four areas of the outcrop are particularly worth viewing. In
area 1, a northeast-trending isoclinal fold plunges gently to the
north. A well-developed cleavage is parallel to its axial surface.
Graded bedding near the hinge of this fold has tentatively been
interpreted to indicate that these folds face downward. Both
limbs of the upright isoclinal fold are cut by cleavage that strikes
northerly and dips steeply; right-hand asymmetrical folds that have
axial surfaces parallel to the second cleavage are found in both
limbs of the upright fold. The concentration of quartz pods marks
a shear zone.

At locality 2, an upright, isoclinal synform deforms an earlier
isoclinal fold. The upright isoclinal synform has an axial sur-
face that strikes northeasterly and dips steeply; its axis plunges
to the northeast. Beds in the hinge of the earlier fold can be
traced around the hinge of the later fold, and cleavage that is
parallel to the axial surface or the earlier fold can be seen only
in its hinge. The plunge of the early fold has not been ascer-
tained. The early fold is thought to relate to the episode of
recumbent folding.

At locality 3, beds of the limestone member are folded by iso-
clinal upright folds. These folds have nearly plane flanks and
sharp hinges. Their axial surfaces strike northeasterly and dip
steeply, and their axes plunge to the northeast. A second
cleavage, best preserved in the beds of phyllite, cuts both
limbs of these folds.

At locality 4, a dike of light-gray plagioclass-quartz-muscovite-
chlorite-calcite granulite strikes N.33°E. and dips 82°SE. It
cuts the bedding at a low angle and has elsewhere been shown to
cut the upright isoclinal folds (Osberg, 1968). A faint cleavage
oriented parallel to the late, north-trending cleavage cuts the
dike. The dike is broken into "rotated" boundinage and the
boudin lines plunge steeply. This dike was intruded after the
formation of the upright isoclinal folds but before the metamor-
phism and late deformation.

Return to cars and continue south on Route 201.

56.7 Stop light. Continue across bridge on Route 201.

56.9 Stop light. Turn left on Route 137.

59.4 Stop 9. Turn left onto Pattee Pond Road and park in vacant lot.
Exposures of Vassalboro Formation form road cut on Route 137.

Interbedded light-gray quartz-mica phyllite and blue-gray,
slightly calcareous quartz wacke are characteristic of the Vassal-
boro Formation. Possible grading and cross lamination may be seen
in the west end of the outcrop.

295

Open and isoclinal upright folds are identifiable in bedding.
Upright folds have axial surfaces that strike northeast and dip
steeply. Their axes plunge to the southwest. The "openness"
of these folds is controlled by the thick quartzose unit exposed
on the south side of the road; away from this unit in less com-
petent lighologies, the folds are more nearly isoclinal. These
folds face upward.

Cleavage is parallel to the axial surfaces of the upright folds.
In competent beds, it is a pressure-solution cleavage; in the
more pelitic beds, it is close-spaced cleavage and micaceous
minerals have formed parallel to the cleavage. The cycloidal
cross sections of thin quartzite beds interbedded with phyllite
may be a pressure-solution effect.

A second set of folds deforms the upright folds and associated
cleavage. These folds have axial surfaces that strike northerly
and dip steeply. Their axes plunge in various directions depend-
ing on the orientations of the surface that they fold. A cleavage
parallels the axial surfaces of these folds.

Return to cars. Reverse direction and proceed west toward Winslow
on Route 137.

61.8 Stop light. Turn left on Route 201 and proceed to Augusta.

77.4 Traffic circle east of business district of Augusta. Follow
traffic circle 240° and bear right on Routes 9 and 17.

77.9 Stop light. Continue straight on Route 9 and proceed to Gardiner.

80.6 Outcrop of Vassalboro Formation on left.

82.8 Stop light. Turn right onto new bridge over Kennebec River. Town
to left is Gardiner.

83.0 Stop light. Turn half-left onto Routes 201, 9, and 126.

83.3 Stop light. Turn right with Routes 126 and 9.

84.2 Bear right with Routes 126 and 9.

85.5 Stop 10. Park at side of road. Be careful to traffic. Exposure
of Cushing Formation and foliated granite.

Biotite-bearing amphibolite and paragneisses consisting of
quartzofeldspathic gneisses and plagioclass gneisses containing
garnet and diopside are intruded by "even"-textured, garnet-bearing
granite gneiss and granitic gneiss containing insets of K feldspar.
Granite and pegmatitic dikes and several orders of veins are pre-
sent. Biotite reaction zones are present along the margins of
several dikes.

296

The amphibolite, paragneiss, and orthogneiss have been deformed
together, and all the rocks have a schistosity that dips modera-
tely toward the northwest. Reclined folds with axial surfaces
that dip more gently than the schistosity deform schistosity as
well as dikes of orthogneiss, pegmatites, and quartz veins. Boudin
lines plunge in approximately the same direction as fold axes.
A fold having an axial surface that strikes northerly and dips
steeply in the western part of the outcrop is interpreted as a
late fold. Enigmatic structures at the middle of the outcrop await
an appropriate explanation.

End of trip.

>

SEP o 6 1989"

DEC 1 7 J993

oversize

QE

78.3

.N4

1980

Edmundston

NEIGC 1 98Q
TRIP COVERAGE

■"-"**—

Fort KeRt

VanBufen

Grand Falls

C-5 S

:/ B-7/C-2

:* <A^M

Caribou

Portage

WashBuch

Ashland

• Presque Isle ^
i ( "i i

B-4 l^3
I fi

Woodstock y

Houlton

-Mount Katahdin

A-2/B-1

/

Patteff 7-

I B-10

A-3/B-2 I • Sheritlftn^x^X i

■Hfrt&*

"linocket

Medway

• Greenville * r

Lincoln

•

A-1

Bangor

urfical Trios Are Underlined

Augusta